shows the Lie de vin Formation (dark, rhythmically banded out-crops) and Igoudine Formation (massive ..... 1953); Adoudounien (CHOUBERT 1956, 1958, 1959); Adou- ...... pleuropsis faunas in Iberia and southern France, has ...... BOUDDA, A., CHOUBERT, G. & FAURE-MURET, A. 1975. Coupe géologique de l'Ounein.
UCL MAGHREB PETROLEUM RESEARCH GROUP (MPRG) Infracambrian/Early Palaeozoic Field Guide Series No. 1
MOROCCO 2006 Ediacaran–Cambrian depositional environments and stratigraphy of the western Atlas regions
Explanatory description and field excursion guide by Gerd Geyer & Ed Landing
Special Issue 6 2006
UCL MAGHREB PETROLEUM RESEARCH GROUP (MPRG) Infracambrian/Early Palaeozoic Field Guide Series No. 1
UCL Maghreb Petroleum Research Group Infracambrian/Early Palaeozoic Field Guide Series
Editorial Board: Dr. Jonathan Craig (Eni, Milan, Italy) apl Prof. Dr. Gerd Geyer (Universität Würzburg, Germany) Prof. Dr. Jürgen Thurow (University College, London, U.K.) Dr. Bindra Thusu (Project Co-ordinator, University College, London,U.K. & Maghreb Petroleum Research Group MPRG)
Special Issue 6 2006
MOROCCO 2006 Ediacaran–Cambrian depositional environments and stratigraphy of the western Atlas regions
Explanatory description and field excursion guide by Gerd Geyer & Ed Landing
COVER ILLUSTRATIONS Front: Tiout section on northern flank of the western Anti-Atlas. This view from top of the Adoudou Formation shows the Lie de vin Formation (dark, rhythmically banded out-crops) and Igoudine Formation (massive beds on nose of mountain). The Souss plain and the southern flank of High Atlas are visible in the background. Low trees are the characteristic arganes. Photo: G. Geyer. Back: Thin section from the lower archaeocyathan bioherm (Ounein A by DEBRENNE & DEBRENNE 1995) in section Le I of the Lemdad Syncline, western High Atlas. The section shows a archaeocyathanmicrobial limestone with the archaeocyathans Archaeopharetra sp. and Chouberticyathus clatratus Debrenne 1964 within a matrix with Epiphython, Girvanella, and Renalcis. Photo E. Berneker.
Beringeria Special Issue 6: 121 S., 69 Abb., Würzburg, 23. November 2006. ISSN 0937-0242
Herausgeber: Freunde der Würzburger Geowissenschaften e.V., Würzburg.
Redaktion: Institut für Paläontologie, Bayerische Julius-Maximilians-Universität, Pleicherwall 1, D-97070 Würzburg.
Druck und Bindearbeiten: Druck + Papier Meyer GmbH, Südring 9, D-91443 Scheinfeld
A contribution to the International Conference
Global Infracambrian Hydrocarbon Systems and the Emerging Potential in North Africa 29 - 30 November, 2006 The Geological Society, Burlington House, London
CONTENTS
G. Geyer & E. Landing Latest Ediacaran and Cambrian of the Moroccan Atlas regions ................... 7
G. Geyer & E. Landing Ediacaran–Cambrian depositional environments and stratigraphy of the western Atlas regions ........................................... 47
References ...................................................................................................... 113
Sketch map of the western and central Anti-Atlas and High Atlas ranges, southern Morocco, illustrating area of Ediacaran and Cambrian rocks (grey), areas which include Cambrian units (light grey), and outcrops of older Neoproterozoic and Mesoproterozoic basement rocks (crosses). Dotted line marks the itinerary of the field excursion from December 2-4, 2006; the number refer to the stops dealt with in the Field Excursion Guide (p. 47-112).
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Latest Ediacaran and Cambrian of the Moroccan Atlas regions GERD GEYER & ED LANDING
Abstract. The characteristics of the latest Ediacaran through Cambrian of the Moroccan Atlas regions are described and illustrated. Such major depositional controls as tectonic environments (a transtensional regime in this time interval) and eustatic changes that defined sequence boundaries and „Grand cycles,“ as well as a progressive southerly movement of the West Gondwanan margin that led to the loss of tropical carbonate platform facies and their replacement by siliciclastic-dominated successions in the Lower–Middle Cambrian boundary interval, are discussed briefly. The lithostratigraphic units for the uppermost Proterozoic(?) to the Upper Cambrian of the Moroccan Anti-Atlas and High Atlas mountains are reviewed and partly revised. The descriptions of lithostratigraphic units include lithology, depositional environments, fossil content, and synonymy. The Jbel Wawrmast Formation is divided into a lower Brèche à Micmacca Member and an upper Tarhoucht Member (new) that comprises the majority of the formation. The bio- and chronostratigraphy of the Atlas regions are summarized, and recently proposed, formal and informal biostratigraphical units of the Lower–Middle Cambrian are reviewed. Detailed stratigraphy allows recognition of diachroneity for several formational contacts. Controversial data and problems of the Moroccan Precambrian–Cambrian boundary are discussed in detail. The available evidence does not permit highly resolved certainty in correlations even at the stage-level with Lower Cambrian sections on other Cambrian continents. However, close similarities exist in the litho- and biostratigraphic developments of southern Morocco and Iberia, and demonstrate that both regions were coterminous on the West Gondwanan margin and geographically separated from the Avalon microcontinent by the latest Proterozoic. Addresses of the authors: Dr. GERD GEYER, Institut für Paläontologie, Bayerische Julius-Maximilians-Universität, Pleicherwall 1, 97070 Würzburg, Germany, e-mail ; Dr. ED LANDING, New York State Geological Survey, New York State Museum, Empire State Plaza, Albany, N.Y. 12230, U.S.A., e-mail .
Introduction The Atlas regions of Morocco have the most complete and best studied latest Ediacaran to Cambrian successions in Africa. These rocks are well exposed and compose the most important reference area for the Lower and lower Middle Cambrian in West Gondwana (GEYER & LANDING 2004). The Anti-Atlas range has comparatively complete, very fossiliferous, and particularly weakly metamorphosed sequences that may extend upward from the uppermost Proterozoic (Fig. 1). Additional Cambrian outcrop areas include the central High Atlas Mountains, the Jbilet and Rehamna regions, and the coastal Meseta, where the rocks were affected by the Hercynian orogeny. In some areas, such as the Rehamna, tectonic activity caused marked metamorphism. In addition, nappe implacement and faulting during the Atlasian phase of the Alpine orogeny created tectonic slices that affected the entire Moroccan part of the Hercynian massif.
This report summarizes the stratigraphic succession of the uppermost Ediacaran and Cambrian of the Moroccan Atlas regions. It follows earlier summaries and stratigraphic revisions by GEYER (1989a, 1990a, 1990b) and GEYER & LANDING (1995), and LANDING et al. (2006), and provides additional details that have led to a somewhat modified view of depositional architecture and controls (Fig. 2). These summaries provided improved definitions for lithostratigraphic units and ended the traditional confusion of litho- and biostratigraphic terms that appeared in reports on the Moroccan uppermost Ediacaran and Cambrian. It might be noted here that the Lower Cambrian in reports by CHOUBERT (1952b) and HUPÉ (1953) began with the lowest, abundantly fossiliferous unit in Morocco, the Amouslek Formation in modern terminology (Fig. 2), and reports after the mid-1970s have greatly extended the lower range of the Cambrian System.
Major depositional controls The late Proterozoic (‘Pan-African’) orogeny in Morocco and northwestern Africa was followed by deposition of Ediacaran and Cambrian cover sequences. The Tiddiline and Ouarzazate Groups are Ediacaran siliciclastic and volcanogenic units, respectively, whereas the Taroudant and Tata Groups (Fig. 2) are Cambrian cover sequences
which reflect an important, long-term transition in lithofacies associations that is well illustrated in the more complete successions of the western and central AntiAtlas. THOMAS et al. (2004) introduced a new lithostratigraphic framework for the Proterozoic rocks of the Anti-Atlas, in which the numerous traditional are partly
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rearranged stratigraphically, partly synoymized, and united into an Anti-Atlas Supergroup and a Ouarzazate Supergroup. Affected are all Ediacaran units which are now grouped in the post-Pan-African Ouarzazate Supergroup, which can be seen as a homonym of the earlier Ouarzazate Group. However, the Ouarzazate Supergroup sensu THOMAS et al. (2004) includes several tectonostratigraphic units which were previously dealt with as independent volcanosedimentary groups or granitoid suites.
SOULAÏMANI et al. 2001, 2003; PIQUÉ 2003). The latter hypothesis is more sound because of the ca. 80 Ma gap between the Pan-African orogeny and the onset of the Tidiline Group deposition. The rifting leading to deposition of the Tidiline Group may be associated with the late Proterozoic breakup of the Rodinia supercontinent, and with the separation and long distance transport of the Avalon microcontinent away from West Gondwana by the latest Proterozoic (LANDING 2005).
Ouarzazate Group Tidiline Group The Tidiline Group (also spelled „Tiddiline“; from Série de Tidiline, LEBLANC 1973) was suggested to be part of a Sarhro Group by THOMAS et al. (2004). The unit consists of fan-glomerates and coarse-grained siliciclastics that were deposited in fault-bounded basin, and rest unconformable on folded Pan-African assemblages and on post-Pan-African granitic intrusives. Conglomerates of the Tidiline Group include clasts from the Bleida granodiorite, for which an age of 579.4 ±1.2 Ma has been reported (DUCROT 1979, INGLIS et al. 2004), and which places a maximum age on the onset of the deposition of Tidiline Group deposition. These rocks have been interpreted as syncollisional molasse (SAQUAQUE et al. 1989; HEFFERAN et al. 1992) and alternatively as rift-related deposits (PIQUÉ et al. 1999,
The traditional Ouarzazate Group (originally termed the ‘Série d’Ouarzazate’) forms the upper part of the Ouarzazate Supergroup in the concept of THOMAS et al. (2004). It consists of an up to ca. 2000 m-thick interval that consists of a wide variety of felsic volcanogenic rocks (e.g., lavas, ignimbrites, and volcaniclastics), and rests on preand synorogenic Pan-African basement. The group is clearly a result of rifting and shows related normal faulting. Volcanic centers of the Ouarzazate Group are intruded by granites and show distinct lateral facies changes, as well as angular discordances (THOMAS et al. 2002). The majority of the Ouarzazate Group was deposited under subaerial conditions. However, lacustrine deposits frequently can be seen to flank volcanic centers and occur within the cores of collapsed calderas.
Fig. 1. Tectonostratigraphic map of the High Atlas and Anti-Atlas regions, Morocco, illustrating areas with pre-Middle Cambrian rocks. Ediacaran and Early Cambrian rocks in bluish colors, older basement rocks in various other colors. Modified from MALOOF et al. (2005), based on SAADI et al. (1985).
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Fig. 2. Cambrian lithostratigraphic units of the Moroccan Atlas regions and their tentative chronostratigraphic distribution. Vertical hachures indicate gaps.
Volcanic rocks of the Ouarzazate Group and associated granites have been dated into a surprisingly narrow time range between 577 and 560 Ma (MIFDAL & PEUCAT 1985; AÏT MALEK et al. 1998; THOMAS et al. 2002; WALSH et al. 2002). Felsic lavas and ash flow tuffs from the uppermost Ouarzazate Group throughout the Anti-Atlas all date between 565 and 560 Ma (MIFDAL & PEUCAT 1985; WALSH et al. 2002; MALOOF 2004). These dates record a rapid cessation of volcanism and onset of thermal subsidence.
Taroudant and Tata Groups: Carbonate-siliciclastic transition Subsidence of the rifted Pan-African margin led to widespread flooding and marine transgression, which initiated the deposition of the Taroudant Group. The Pan-African cover sequence represented by the Taroudant and Tata Groups includes a thick interval of carbonate-dominated rocks followed by siliciclastic-dominated rocks higher in the Cambrian in the western AntiAtlas. The uppermost Ediacaran–middle Lower Cambrian in this region (i.e., the Adoudou, Lie de vin, and Igoudine Formations) is dominated by restricted marine, inner platform limestones and dolostones and is overlain by a shale–carbonate ‘Grand cycle’ in the upper Lower Cambrian (i.e., the Amouslek Formation). Uppermost Lower Cambrian and lowest Middle Cambrian units include finegrained siliciclastics with relatively minor nodular and
bedded limestones and very minor volcanic ashes (Issafen Formation) overlain by higher-energy, latest Lower to lower Middle Cambrian, shallow marine, sandstonedominated facies (i.e., the Tatelt and Tazlaft Formations). Limited carbonates also characterize the lower Middle Cambrian, and fossil-hash limestone beds are regularly encountered only in the lower part of the Jbel Wawrmast Formation (i.e., the Brèche a Micmacca Member in Fig. 2). THEOKRITOFF’s (1979) early suggestion that Morocco came under progressively stronger, higher latitude climatic influences during the Cambrian seems to be appropriate in explaining the decrease in limestones from the latest Ediacaran–Lower Cambrian, the loss of archaeocyathan build-ups higher in the Lower Cambrian, and in development of paleoenvironments that permitted trilobite to immigrate in the Lower–Middle Cambrian boundary interval between the separate continents of Gondwana and Avalon (LANDING 1996). (It should be noted, however, that the appearance of special trilobite clades may indicate that Morocco plays a key role for the diversification of trilobites in ‘Acadobaltica.’) A rapid motion of the Moroccan margin of Gondwana seems to have taken place from the low latitude position (BURRETT et al. 1990) suggested by latest Proterozoic– Lower Cambrian carbonate facies to the high south temperate latitudes indicated by late Early and Middle
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Cambrian paleomagnetic work (e.g., SCOTESE et al. 1979; SMITH et al. 1981).
Tectonic setting The presence of volcanic rocks and a volcanoclastic component of sedimentary rocks has long been recognized in the uppermost Ediacaran–Middle Cambrian cover sequence of Morocco (e.g., DESTOMBES et al. 1985; BUGGISCH & FLÜGEL 1988; BUGGISCH & SIEGERT 1988). These volcanic rocks indicate that deposition of the shallow water cover sequence did not take place on a completely ‘passive margin.’ However, the relative amount of volcanic rocks is much less in the cover sequence than in parts of the underlying Pan-African orogen, and different tectonic environments are represented. Two important intervals of volcanism are recorded in the Taroudant and Tata Group cover sequence. These volcanic intervals correspond to basin reorganizations. Volcanic ashes and flows are interbedded locally in the lower onlap dolostones of the Adoudou Formation in the lowest part of the cover sequence (e.g., CHOUBERT & FAURE-MURET 1970). One source for these ashes is dramatically preserved as the „Volcano of Alougoum“ in the eastern Graara massif (CHOUBERT 1952a; CHOUBERT et al. 1979). An early U/Pb date of 534 ± 10 Ma from the Volcano of Alougoum (DUCROT et al. 1976, recalculated by DUCROT & LANCELOT 1977) suggests that deposition of the Adoudou Formation began in the middle part of the sub-trilobitic Lower Cambrian (i.e., middle Placentian Series and Nemakit-Daldynian equivalent; see ISACHSEN et al. 1994; LANDING 1994, 1995a) and that the Precambrian–Cambrian boundary in this part of the Anti-Atlas may lie at the unconformity between the Pan-African orogen and overlying Adoudou Formation. The temporal relationship of volcanism and marine onlap holds not only in the earliest Cambrian but also in the Lower–Middle Cambrian boundary interval. BUGGISCH & SIEGERT (1988) proposed a northwestern source of volcanoclastic debris in the upper Lower Cambrian of the Anti-Atlas. VILAND (1972), BOUDDA et al. (1972), DESTOMBES et al. (1985), and our field work indicate that volcanoclastic sandstones, thin K-bentonites and local pillowed flows are regularly encountered in the lowest Middle Cambrian Tatelt Formation and overlying lower part of the Jbel Wawrmast Formation of the central and western Anti-Atlas, the Jbel Sarhro, and the High Atlas. This volcanism accompanied uplift and erosion of blocks in the central High Atlas, downdropping of the Souss Basin in the Anti-Atlas (discussed below), and unconformable onlap of the lower Middle Cambrian across the lower Lower Cambrian on the south side of the Jbel Sarhro and across the Pan-African orogen on the north
side of the Jbel Sarhro and the Jbel Ougnate in the eastern Anti-Atlas. The association of lower Lower and lower Middle Cambrian onlap with volcanism suggests that epeirogenic processes in a pull-apart (strike-slip or transtensional) regime may have defined the Cambrian depositional basin in Morocco. Minor volcanism is recorded through most of the Cambrian of Morocco, but other volcanic intervals do not seem to have a relationship to basin reorganization. For example, COMPSTON et al. (1990, 1992) determined a SHRIMP date (discussed below) on a thin ash in the higher Lie de vin Formation. Our stratigraphic reevaluation shows that the thin volcanic debris layers of the „Série de Jebel Tichinchine“ (designation abandoned, GEYER & LANDING 1995) reported by BUGGISCH & SIEGERT (1988) actually occur in the upper Lower Cambrian Issafen Formation. Ashes in the Lemdad Formation will be seen in the course of the field trip (see Stops 4 and 7, field excursion guide, this volume). Beds of volcanic ash in the upper Middle Cambrian Tabanite Group are also noted in maps of the Jbel Sarhro area published by the Service Géologique du Maroc. However, these horizons have been determined to be non-volcanogenic arkoses (E. LANDING and W. HELDMAIER, unpubl. data).
Souss Basin The uppermost Proterozoic–Cambrian cover sequence of the Anti-Atlas and High Atlas mountains was deposited in a sedimentary basin that persisted into the Silurian. This Souss Basin (GEYER 1989a) is generally similar to an epicratonic basin, with exception of the continuing volcanism that persisted into late Middle Cambrian times. The axis of the basin roughly coincides with the modern trend of the Anti-Atlas (Fig. 1). The depocenter of the Souss Basin obviously underwent a counter-clockwise reorientation from a northwestern trend during Adoudou Formation deposition to a WSW-trend during deposition of the Middle Cambrian Jbel Afraou Formation and, possibly, the younger formations of the Tabanite Group. The lower part of the cover sequence has been divided into groups that reflect more-or-less asymmetrical transgressive-regressive cycles. Depositional environments changed through time with shoreline offlap and the increasing influx of detrital material derived from the east. Biostratigraphic evidence for diachroneity in formation contacts also suggests that onlap-offlap cycles are recognizable in the Cambrian of Morocco (Fig. 2). However, formational contacts are traditionally defined through seemingly gradational intervals in the Moroccan Cambrian.
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Cambrian stratigraphy Taroudant Group The lowest part of the Moroccan cover sequence is known from the western and central Anti-Atlas and from the western and, probably, the eastern margins of the Adrar n’Dren massif of the central High Atlas. The Taroudant Group (GEYER 1989a) rests unconformably on metamorphic and volcanic complexes of the late Proterozoic Pan-African orogen. Vast areas of the western AntiAtlas are underlain by the Taroudant Group, and a distinctive topography is developed on the group. The Taroudant Group has received special interest because many workers have assumed that the Precambrian–Cambrian boundary is located within it. The available biostratigraphic data is limited on the Taroudant Group because most of its facies represent proximal, restricted (frequently hypersaline) environments that precluded colonization by animals that produced ichno-
fossils useful in latest Ediacaran–earliest Cambrian biostratigraphy. Similarly, although the diversification of Early Cambrian skeletalized metazoans took place in peritidal habitats (LANDING & WESTROP 2004), the nearshore habitats of the Taroudant Group were generally too hypersaline for colonization by these animals. However, resolution in Lower Cambrian correlations is provided by carbon isotope, magnetostratigraphic, and geochronologic investigations. The basic assumption in these studies that non-conventional correlation techniques can be used for interregional correlation is problematical because these techniques cannot be related to an adequate biostratigraphic standard in the Taroudant Group. However, a 534 ± 10 Ma date on the lower Adoudou Formation and revised interpretations of chemostratigraphic data (MALOOF et al. 2005) suggest that almost all of the Taroudant Group is Lower Cambrian. In addition,
Fig. 3. Sketch map of the High Atlas and Anti-Atlas regions, Morocco, illustrating the distribution of the Adoudou Formation and the tentative paleogeography during its deposition. The line that starts at Agadir indicates a zone of thrust faults, which separate the Anti-Atlas domain from the Hercynian fold belt. Coarse stipples indicate the probable limits of the distribution of the Tabia Member. Irregular pattern represents non-depositional areas; small stipples = marginal clastic belt; irregularly hachured = areas with significant content of volcanic rocks; isopach lines indicate probable thicknesses (in meters). The white basinal center marks the major development of carbonates and the thickest parts of the formation, and is consistent with synsedimentary halfgraben structures and pronounced subsidence as described by BENSAOU & HAMOUMI (2001, 2003). Modified from GEYER (1989a: Fig. 2), based on data from G. CHOUBERT and unpublished data from G. GEYER.
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our work with S. A. BOWRING on U/Pb volcanic zircon dates on the underlying, upper part of the Ouarzazate Group, the lower Adoudou Formation, and middle Lie de vin Formation are expected to provide additional precise dates to bracket the Pan-African orogen–lower Taroudant Group unconformity and determine approximate rates of lowest Cambrian accumulation in the Anti-Atlas ranges. The Taroudant Group is of interest because of its lithofacies associations. The group includes a depositional cycle developed on a shallow marine carbonate platform (Adoudou Formation) and the higher carbonate-shale sequence (Lie de vin Formation) and is the best preserved, seemingly continuous sequence through lowest Cambrian, extremely shallow marine environments.
west are replaced by a more siliciclastic-rich facies in the central Anti-Atlas. The change in facies is generally a lateral transition rather than an interfingering of the different facies. Although the Adoudou Formation is one of the thickest and best exposed units in Morocco, detailed information about its lateral and vertical facies changes and the continuity of its deposition is limited. Detailed studies are only available on the Tiout section (MONNINGER 1979; see Stop 1, field excursion guide, this volume) and, to a more limited extent, for the Tamjout section (CHAZAN 1954). A more-or-less detailed investigation on the carbon isotope signatures of the Adoudou Formation was published by MAGARITZ et al. (1991) that they claimed permitted a comparison with the Siberian Yudoma and Pestrotsvet Formations (see Precambrian–Cambrian boundary discussion, below).
Adoudou Formation Depositional environment Lithology and regional development The Adoudou Formation (HOLLARD 1985, named after Oued Adoudou in the western Anti-Atlas, has a relatively thin (up to 250 m), lower unit composed of conglomeratic, dolomitic, and siliciclastic beds with local volcanic ashes and flows. This lower interval of the Adoudou Formation was recently named the Tabia Member by MALOOF et al. (2005). The name should not be confused with the „Série d’Igmir Tabia“ which is a formation composed of shales and carbonates and forms the terminal unit of the Précambrien III in the Ouaoufenrha massif. The Tabia Member is overlain by limestones and dolostones of the Tifnout Member (MALOOF et al. 2005; earlier termed the ‘Limestone Member’ by GEYER, 1989a). The formation forms a broad outcrop belt in the core of the western and central Anti-Atlas where it is primarily responsible for the topography. However, the formation wedges out at the margins of the numerous Proterozoic inliers to the east (BOUDDA et al. 1979; VAN LOOY 1985). These regional variations and the resulting facies changes led BENSAOU & HAMOUMI (2001) to suggest numerous members, based on their studies at the northeastern rim of the Kerdous massif and adjacent areas. In fact, these „members“ represent lithofacies belts rather than lithostratigraphic units. In addition, the Adoudou Formation is known from the south-central High Atlas Mountains (e.g., the Lemdad syncline and Agoundis sections). The thickness of the Tifnout Member ranges from 500 to more than 1,000 m. Apart from the local and regional facies changes, a west-to-east lithofacies transition takes place in the Adoudou Formation. Relatively pure carbonates in the
Adoudou Formation carbonates reflect deposition primarily under shallow subtidal to inter- and peritidal marine conditions as a result of thermal subsidence. The blanket of dolostones covers virtually the entire western and central Anti-Atlas. Topographic differences during the deposition resulted in dramatic lateral facies changes as demonstrated by BENSAOU & HAMOUMI (2001). Deposition of the Tabia Member took place in a series of linked half-grabens bounded by steep normal faults (BENSAOU & HAMOUMI 1999a, 1999b, 2001, 2003). This development is nicely recorded by differences in thickness and facies and is easily recognizable on paleogeographic maps (see Fig. 3). An interesting lithofacies association includes the interbedded rhyolitic or syenitic explosive volcanics and planar-laminated, stromatolitic, pinkish-brown dolostones with scattered anhydrite molds at the western end of the El Graara massif (i.e., BUGGISCH & FLÜGEL 1986; LANDING & HELDMAIER, unpub. data). CHOUBERT (1952a) documented intercalations of trachyandesitic rocks in the formation. The volcanic source at Jbel Bokho (Jebel Boho) near Alougoum (a syenite plug dated at 534 ± 10 Ma; DUCROT & LANCELOT 1977) is still recognizable as a volcano (CHOUBERT 1952a; BOUDDA et al. 1979), and the middle of the Adoudou Formation in the surrounding area is dominated by volcanic rock. Onlap of the formation took place across an erosion surface on the late Proterozoic Pan-African orogen. The transition from the Tabia Member into the Tifnout Member reflects a change from a more siliciclastic onlap sequence into a predominantly carbonate sequence with probable shoreline transgression.
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Fig. 4. Adoudou Formation, typical development of massive thickbedded carbonates of the Tifnout Member, near Asdrâm Amrazi.
Fossils Reports on Adoudou Formation fossils are limited. CHOUBERT et al. (1950, 1952a, 1952b, 1987), BUGGISCH et al. (1978), BUGGISCH & HEINITZ (1984), and VAN LOOY (1985) reported stromatolites from the Tifnout Member in the Anti-Atlas. The presence of a skeletal alga referred to Kundatia composita has been described from the central Anti-Atlas (BUGGISCH & FLÜGEL 1986) although this determination must be regarded as erroneous (see below under „Moroccan Precambrian–Cambrian boundary“). CHOUBERT et al. (1979) reported a number of acritarchs from the Tabia Member at localities in the western AntiAtlas. HOUZAY (1979) described discoid structures from tidally dominated shales of the uppermost Tabia Member in the Igherm area and interpreted them as Ediacarantype fossils. However, the Ediacaran nature of these structures could not be proved by subsequent studies. In addition, ribbon-like structures similar to vendotaenids were found in green shales of the Ouaoufenrha area on the northern slope of the western Anti-Atlas (G. Geyer, unpubl. data; see Field Guide section, Stop 9).
Tabia Member
Synonyms
The Tifnout Member (MALOOF et al. 2005; Limestone Member of GEYER 1989a; Calcaires inférieures s.s.) is a rather monotonous sequence of shallow marine carbonates that reaches about 1,000 m in the western AntiAtlas (Oued Sdas section) but only 200 m in the Bou Azzer area of the central Anti-Atlas (El Graara massif). Outcrops in the central High Atlas are composed of massive, light-grey dolostones that lack distinct bedding. The light colored, massive dolostone and subordinate limestone of the member are interbedded in most sections with mudstones, shales, and sandstones. Meterscale, shallowing-upward parasequences are developed where the thickness is not exceedingly high. Such parasequences begin with variably colored siltstones and
Calcaires inférieurs (CHOUBERT 1952b, CHOUBERT in HUPÉ 1953); „Série de base“ plus „Calcaires inférieurs“ (CHOUBERT in HUPÉ 1953); Adoudounien (CHOUBERT 1956, 1958, 1959); Adoudounien inférieur (CHOUBERT 1963, CHOUBERT et. al. 1975); Dolomies et Calcaires inférieurs (BOUDDA & CHOUBERT 1972, CHOUBERT et al. 1975); Série de base plus Calcaires et Dolomies inférieurs (BOUDDA et al. 1979); Adoudounien p. p. (BOUDDA et al. 1979); Adoudounien s. s. (CHOUBERT & FAURE-MURET 1987); Lower Limestones and Dolomites Formation (HOLLARD 1985); Série des calcaires inférieurs (BUGGISCH & FLÜGEL 1988); Adoudou Formation (GEYER 1989a); Adoudon Formation (MAGARITZ et al. 1991). The term „Calcaires inférieurs“ has been misspelled several ways in relatively recent publications in English language.
The Tabia Member (Série de base; CHOUBERT 1952b; Basal Member, HOLLARD 1985 and GEYER 1989a) is an approximately 100 m-thick (0-250 m), transgressive unit. Three units are distinguished in completely developed sections of the Tabia Member in the western Anti-Atlas: (1) lower, massive conglomerates, (2) middle, diagenetically recrystallized dolostones and limestones with mudcracked siltstones and evaporites, and (3) an upper, shale-dominated interval with conglomeratic, quartz arenite, or arkose beds and limestone and dolostone layers (BOUDDA et al. 1979; HOLLARD 1985). Conglomerates replace the carbonates towards the eastern margin of the basin. The Tabia Member is not known with certainty east of the Jbel Sirwa massif. A facies change into siliciclastic rocks takes place west of the Adrar n’Dren massif in the Agoundis area of the High Atlas, where VILAND (1972) noted the association of green shales with subordinate sandstones and carbonates. Tifnout Member
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marlstones, followed by laminated micritic limestones and dolostones, respectively, more seldom by grainstones, and are capped by laminated carbonates with microbial mats and sometimes stromatolites. Such carbonates often show ripple marks and reworked argillaceous clasts, and local signs of subaerial exposure (tepee structures, desiccation cracks, pseudomorphs after gypsum, etc.; MONNINGER 1979). Rheological instability led to common slump folding, brecciation, and other synsedimentary deformational structures (BUGGISCH & HEINITZ 1984; HEINITZ 1984). The prominent dolomitization, particularly in the lower and middle part of the Adoudou Formation, is striking by comparison with overlying formations and has been linked to an increased Mg/Ca ratio or a higher partial pressure of CO2 in the Precambrian by TUCKER (1992). This explanation for Adoudou dolostones is considered unlikely; the lower Adoudou Formation is lowest Cambrian on the basis of available geochronologic data, and the dominant carbonate mineral in earliest Cambrian sediments elsewhere has been noted to be aragonite (LANDING 1992). The dolomitic sequence of the Adoudou Formation may simply reflect highly restricted inner platform deposition or even burial dolomitization by
hydrothermal fluids during the Hercynian or Alpine orogeny. CHAZAN (1954) described a lower submember of the Tifnout Member („Assise de Tamjour“) characterized by silicified dolostones with isolated chert lenses. The siliceous intercalations or silicified carbonates are often variably mineralized. Lie de vin Formation Lithology and regional extent The Lie de vin Formation (HOLLARD 1985) consists of fine-grained siliciclastic rocks with numerous limestone and dolostone beds that locally result in conspicuous color-banding. The term lie de vin (CHOUBERT 1942) refers to the typical purple or burgundy color of the shales in the unit and not to a geographic place and type section, which actually should be the basis for the name of any lithostratigraphic unit. The known thickness ranges from ca. 300 m at Warzazate (BOUDDA et al. 1979) to about 950 m in the type section at Tiout (MONNINGER 1979). These measurements illustrate the easterly thinning of the unit to about 25 m at the southeastern rim of the Jbel Sarhro (DESTOMBES, unpublished data)
Fig. 5. Sketch map of the High Atlas and Anti-Atlas regions, Morocco, illustrating the distribution of the Lie de vin Formation and the tentative paleogeography during its deposition. The line that starts at Agadir indicates a zone of thrust faults, which separate the Anti-Atlas domain from the Hercynian fold belt. Irregular pattern depicts non-depositional areas; small stipples = marginal siliciclastic belt; brick signature = areas with significant to dominant carbonate deposits (consistent with a synsedimentary half-graben structure and pronounced subsidence); isopach lines indicate suposed thicknesses (in meters). Modified from GEYER (1989a: Fig. 3) and based on data from G. CHOUBERT and unpublished data from G. GEYER.
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Fig. 6. Lie de vin Formation at the eastern flank of the Issafen Syncline, view towards south illustrates the dipping towards the core of the syncline and the name-giving reddish color and the typical banding due to numerous obvious parasequences.
and less than 15 m at the El Graara massif (MALOOF et al. 2005). The Lie de vin Formation usually overlies the Adoudou Formation, but it locally rests unconformably on the Pan-African basement east of the Sirwa dome. The rocks of the Lie de vin Formation vary in the relative proportion of carbonates and siliciclastics. The increase in siliciclastics over the Adoudou Formation reflects a relative regression. Red sandstones and conglomerates dominate at the basin margin in the east. In the central Anti-Atlas, the middle part of the formation is a completely detrital succession dominated by sandstones that interfinger with finer-grained sediments. This lithofacies association progrades from the east of Warzazate into the central Anti-Atlas. This siliciclastic interval is the Tikirt Member (from Grès de Tikirt; GENTIL 1912) that crops out from the west-central Anti-Atlas to the Jbel Sarhro in the east. Characteristic Lie de vin Formation sections are found on the southern slope of the western Anti-Atlas, where sections are dominated by purple shales. Dolostones and limestones become increasingly common towards the top of the formation, especially in the upper third of the sequence (HOLLARD 1985). The type section at Tiout is somewhat atypical in the dominance by massive, brownish weathering limestones and dolostones with relatively thin siltstones.
Depositional environment Seaward tilting of the Anti-Atlas margin at the end of the Adoudou deposition initiated a change in deposition marked by the Lie de vin Formation. This tilting, or flexure is recorded by a significant landward regression and resultant erosional contact of the Lie de vin Formation with underlying units and deposition of fluvial siliciclastics (Tikirt Member; e.g., CHBANI et al. 1999). Seaward, the formation shows a change from peritidal carbonate to sublittoral shale-biohermal limestone sedimentation with an increase of terrigenous detritus. The Tiout section shows a complex interbedding of lithofacies that composes ca. 180 of autocyclic shallowing-upward parasequences (see MONNINGER 1979; TUCKER 1986b; GEYER 1989b). Complete parasequences start on a flooding surface followed by interbedded argillites and laminated carbonates, which are then overlain by stromatolitic limestones and/or thrombolitic carbonates. Biohermal carbonates pass laterally into grainstones. Dolomitic, limy, marly, and silty sediment were deposited under persistant low energy, shallow, relatively hypersaline conditions (MONNINGER 1979). The alternating lithofacies reflect some combination of changes in siliciclastic input, salinity, or climate or small-scale sea level changes. These conditions led to symmetrical cycles (or alternations) that include tidal flat dolostones
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Fig. 7. Chemostratigraphic analysis of carbon isotopic signatures in eight key sections of Ediacaran–Cambrian units in the western Anti-Atlas. Lithofacies are strongly simplified and only show dominating lithofacies the sections. All isotopic data from MALOOF et al. (2005) except from Tiout (from TUCKER 1986 and MARGARITZ et al. 1991). Modified from MALOOF et al. (2005: Fig. 2).
with cyanobacterial laminates and such indicators of hypersaline conditions as pseudomorphs after halite and anhydrite or gypsum (MONNINGER 1979). The evaporitic facies pass offshore into micritic, peloidal, and oolitic limestone with stromatolite and thrombolite bioherms (SCHMITT 1979a; TUCKER 1986a). The lower part of the Lie de vin Formation, in particular, has facies that suggest a non-marine origin (LATHAM 1990). The fluvial Tikirt Member (Fig. 8) in the middle Lie de vin Formation is a coarser, siliciclastic facies that reflects regression and terrigenous input. The lower part of the Tikirt has flute and load casts (BUGGISCH et al. 1978). Fossils The Lie de vin Formation is locally fossiliferous, but animal body fossils are not yet known. Stromatolites and thrombolites are locally conspicuous and form large bioherms in the Anti-Atlas. They were studied in detail at
Tiout and include Linella, Tungussia, Patomia, Nimbophyton, and Igoudinia (SCHMITT 1979a, 1979b). However, their biostratigraphic interpretation has generated controversy (BERTRAND-SAFATI 1981; compare CHOUBERT & FAURE-MURET 1987, and SDZUY & GEYER 1988) and has not resolved the problem of correlation. Thrombolite construction is related to the occurrence of the probable cyanobacteria Tarthinia, Renalcis, and Kordephyton, an assemblage that suggests an age equivalent with the Siberian Tommotian Stage or younger (LATHAM & RIDING 1990). The upper part of the formation in the western AntiAtlas has complex metazoan traces such as Diplocraterion (LATHAM & RIDING 1990), which are best interpreted as indicating a Cambrian age. Trilobites which were claimed to be found at the top of the Lie de vin Formation in the Lemdad River section in the central High Atlas (BOUDDA & CHOUBERT 1972)
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Fig. 8. Spectacular, cliff-forming sandstones of the Tikirt Member in the canyon southeast of Tizi n’Tinififft, southeast of Warzazate.
were reinterpreted as derived from the base of the Lemdad Formation (SDZUY 1978; assigned to the „Calcaire Supérieur“; see discussion under Stop 4, Field Excursion Guide, this volume). Trace fossils are also in the upper part of the Tikirt Member. These include Diplichnites-type traces. Synonyms and approximately corresponding terms Série régressive lie-de-vin (CHOUBERT 1952b); Série régressive des schistes lie de vin (CHOUBERT 1952b); Schistes lie de vin (CHOUBERT 1952b); Série regres-sive lie de vin (CHOUBERT in HUPÉ 1953); Géorgien inférieur p. p. (CHOUBERT 1956, 1958); Adoudounien supérieur p. p. (CHOUBERT 1963); Adoudounien moyen (CHOUBERT & FAURE-MURET 1970); Série lie de vin (CHOUBERT 1963, BOUDDA & CHOUBERT 1972, CHOUBERT et al. 1975); Adoudounien supérieur (BOUDDA & CHOUBERT 1972, CHOUBERT et al. 1975); Taliwinien (CHOUBERT et al. 1975, BOUDDA et al. 1979); Lie de vin Formation (GEYER 1989a, 1989b).
Tata Group The Tata Group (GEYER, 1989a) crops out in the western and central Anti-Atlas and in the bloc occidental of the Adrar n’Dren massif, central High Atlas. The group composes the remainder of the Lower Cambrian above the Lie de vin Formation. The recognition of the Hupeolenus Zone from the Tatelt Formation and the very top of the underlying Issafen Formation (e.g., GEYER 1990a, LANDING et al. 2006) means that the Tata Group extends into the lowest Middle Cambrian (lowest Agdzian Stage) of Morocco. The alternative explanation that th abrupt faunal change from the Moroccan upper Lower Cambrian Sectigena Zone into the Hupeolenus Zone corresponds to a cryptic, but major, sequence boundary and unconformity within the former „Asrir Formation“ was shown to be incorrect due to fossil evidence from the Issafen Formation
of the Lemdad syncline (GEYER 1983, GEYER & LANDING 1995, LANDING et al. 2006; see Stops 5–7, field excursion guide, this volume). Tata Group outcrops cover large areas and are seen in the well developed sections in the western Anti-Atlas. This region is the classic Lower Cambrian field region for such pioneers of the Moroccan Cambrian as the late Georges Choubert and Pierre Hupé (e.g., HUPÉ 1952, 1953; CHOUBERT 1953). Igoudine Formation Lithology and regional development The Igoudine Formation (GEYER, 1989a) is composed of massive limestone and dolostone beds (MONNINGER 1979; GEYER 1989a). It forms large outcrops along the border of the western and central Anti-Atlas. The thickness varies regionally and reaches almost 400 m in the Tiznit area. Deposition of the Igoudine Formation was initiated by regional deepening from the restricted marine interval of the underlying Lie de vin Formation. The base of the formation is usually sharp and marked by the appearance of massive carbonates with thin shales. The formation consists mainly of dark dolostones and limestones and subordinate light grey to pink carbonates. Conspicuous regional and vertical changes are present in the Igoudine Formation. The formation is composed of massive carbonates that are largely dolomitic at the base and become increasingly calcareous upward in the southern Anti-Atlas. Red, pink, or light colored cavernous limestones and brown marlstones are intercalated in the formation in this region (HOLLARD 1985), which indicate meter-scale parasequences. However, shales and sandstones with ripple marks, cross-beds, intraclasts,
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and reworked, low-amplitude cyanobacterial mats are more frequent on the southern flank of the Jbel Sarhro (BUGGISCH et al. 1978) and indicate higher energy conditions locally in the upper part of the Igoudine Formation. Synsedimentary deformations are conspicuous in the vicinity of the Jbel Sirwa (see BUGGISCH & HEINITZ 1984, and HEINITZ 1984, for analysis). An important lithofacies change can be identified in several sections. This includes a change from black, often laminated, lime mudstones and dolosiltites into overlying, massive, dark ooid pack- and grainstones with intraclasts and marlstones with thin quartz sandstone interbeds. The overlying lithofacies is the Tiout Member (GEYER, 1989a; i.e., „Transition layers“ sensu MONNINGER 1979, and SCHMITT 1979a; possibly identical with the „Calcaires de base“ sensu BUGGISCH et al. 1978). The amply fossiliferous Tiout Member has a thickness of about 110 m at Tiout (see detailed description in the Field Trip Guide, Stop 1, this volume), slightly more than 50 m in the Amouslek region, and about 20 m south of Jbel Kissane.
Depositional environment The Igoudine Formation was deposited under shallow marine conditions, locally under restricted circulation and relatively high salinity as indicated by frequent low algal mats and stromatolitic buildups. BUGGISCH et al. (1978) noted dessication cracks and ignimbrite horizons as evidence for subaerial exposure. Higher energy conditions, possibly associated with slightly deeper and less restricted marine conditions, are recorded by the Tiout Member (MONNINGER 1979). The development of the Tiout Member facies may mark a transition from nearshore to open shelf conditions. The appearance of a skeletalized fauna (archaeocyathans, trilobites, hyoliths, chancelloriids) in the Tiout Member is probably linked to these facies changes. Thin volcanic ashes appear in the Tiout Member, and are under analysis by S. A. Bowring for U–Pb dates. Fossils The low energy facies of the lower part of the forma-tion are frequently stromatolitic. Reports on these build-ups
Fig. 9. Sketch map of the High Atlas and Anti-Atlas regions, Morocco, illustrating the distribution of the Igoudine, Amouslek, Lemdad, and Tislit Formations and the tentative paleogeography during their deposition. The line that starts at Agadir indicates a zone of thrust faults, which separate the Anti-Atlas domain from the Hercynian fold belt. Irregular pattern represents nondepositional areas; large stipples = marginal clastic belt; small stipples = tentative distribution of Tislit Formation; spotted brick signature = tentative distribution of Lemdad Formation; brick signature = areas with considerable to dominant carbonate deposits in the Amouslek Formation (linked with shoals in the basin and resulting reef developments). Based on maps from GEYER (1989a: Fig. 4) and CHOUBERT and unpublished data of G. GEYER.
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Fig. 10. Carbonate-dominated Igoudine Formation (with minor bioconstructions left of the center of the photo and to the lower right) overlain by shale-carbonate parasequences of the Tislit Formation. Valley close to Talate, west of the El Graara Massif.
are based on sections near the Jbel Sirwa massif (BUGGISCH & HEINITZ 1984; HEINITZ 1984), and biostromal and biohermal stromatolites are also known from the southern Jbel Sarhro (BUGGISCH et al. 1978). The most comprehensive study on these stromatolites deals with the lower Igoudine Formation at Tiout (SCHMITT 1978, 1979a, 1979b), where trace fossils also occur. The stromatolites Acaciella sp. cf. A. angepena and the endemic Tioutella bouddai (SCHMITT 1979a) occur at Tiout. Horizons with thrombolites are also known from the lower member of the Igoudine Formation at Tiout (SCHMITT & MONNINGER 1977; SCHMITT 1979a). Studies on trace fossils have not been done. The Tiout Member has the oldest known skeletal fossils known from Morocco. They include trilobites, archaeocyathans, chancellorids, hyoliths, Hyolithellus, Coleoloides, and the calcimicrobes (including prob-able cyanobacteria) Renalcis, Epiphyton, Tolypothrix, Girvanella, and others (SDZUY 1978; DEBRENNE & DEBRENNE 1978, 1995; and unpubl. information of E. Berneker and G. Geyer). The trilobites include species of Eofallotaspis, Hupetina antiqua, and yet undescribed bigotinoid redlichiaceans of the Eofallotaspis and Fallotaspis tazemmourtensis Zone (SDZUY 1978, 1981; GEYER 1996). Archaeocyathans form bioherms in some beds of the Tiout Member (DEBRENNE 1960, 1975; DEBRENNE & DEBRENNE 1976, 1978). Similar faunas are known to occur at Tazemmourt and Amouslek (SDZUY 1978; DEBRENNE 1960; DEBRENNE & DEBRENNE 1995). The archaeocyathans from the Tiout section belong to an Erismacoscinus fasciolus–Retecoscinus minutus Zone (DEBRENNE & DEBRENNE 1978, 1995) and have been used for intercontinental correlation because they were thought to indicate an age coeval to that of the upper Atdabanian stage of the Siberian Platform (ROZANOV & DEBRENNE
1974, MAGARITZ et al. 1991). However, regional correlation inside Morocco and correlation by trilobites into other regions prove that this correlation of the Tiout Member is erroneous and that the Tiout Member is older that the upper Atdabanian in Siberia. Synonyms Calcaires supérieurs (CHOUBERT 1952b, CHOUBERT in HUPÉ 1953); Géorgien inférieur p. p. (CHOUBERT 1956, 1958); Calcaires géorgiens (CHOUBERT 1958); Adoudounien supérieur p. p. (CHOUBERT 1963); Adoudounien supérieur (CHOUBERT & F AURE-M URET 1970); Calcaires supérieurs avec Collenia (BOUDDA & CHOUBERT 1972, CHOUBERT et al. 1975); Ouneinien (BOUDDA & CHOUBERT 1972); Étage Ouneinien (CHOUBERT et al. 1975); Upper limestone formation p. p. (HOLLARD 1985); Upper Limestones Formation p. p. (HOLLARD 1985); Série des calcaires supérieurs + Calcaires de base (BUGGISCH & FLÜGEL 1988); Igoudine Formation (GEYER 1989a, 1989b, 1990b).
Amouslek Formation Lithology and regional development The lowest part of the classical Lower Cambrian (i.e., Georgien) of HUPÉ (1953) includes an alternation of generally dark limestones and lighter siliciclastic mud-stones first termed the „Série schisto-calcaire“ (CHOUBERT 1952b) and now known as the Amouslek Formation (GEYER 1989a, 1990b). The „Amouslekien“ of CHOUBERT et al. (1975) is not identical to HUPÉ’s (1960) „Etage d’Amouslek.“ The latter is a biostratigraphic unit and differs from the diachronous „Série schisto-calcaire.“ The Amouslek Formation is a sequence of greenish, purple, or grey shales (often slaty) with intercalated limestones, the latter often containing archaeocyathan bioherms. The unit is widely distributed in the Anti-Atlas and ranges from about 220 m in thickness at the Amouslek type section to about 40 m south of Jbel Kissane. It
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wedges out towards the eastern part of the Jbel Sarhro (BOUDDA et al. 1979). The amount of siliciclastic content increases in the shallower water environments to the east (CHOUBERT 1964) with a lateral transition into the Tislit Formation in the Tazenakht area. Volcanic material is present at some localities, as at Tazemmourt (SCHAER in BOUDDA et al. 1979), but significant volcanic activity is not recorded by the formation. The transition from the Igoudine into the Amouslek Formation is an abrupt change from massive carbonates to mixed shale/slate–carbonate interbeds. In the western Anti-Atlas section at Tiout, a lensing, thin volcaniclastic sandstone bed with reworked clasts of Igoudine carbonate suggests an unconformity and sequence boundary at the base of the Amouslek Formation (E. LANDING & G. GEYER, unpub. data) A typical lithologic sequence is exposed at the classical locality of Amouslek on the northern slope of the western Anti-Atlas. This type section of the Amouslek Formation consists of: (1) lower green slates, (2) a middle unit of grey sandy or red muddy slates, and (3) a thick upper unit of limestones with green and light yellow argillaceous shales. Shale–oolite alternations occur in this part of the section in the northwestern Anti-Atlas. These asymmetric cycles/alternations commence with olive shales that grade up into quartz arenites and oolitic limestones with desiccation (or synaeresis?) cracks and symmetrical and interference ripples and finally into cross-bedded oolites (TUCKER 1986b). The tops of these alternations locally have thrombolite build-ups. The increase and dominance of shoal-water limestones upward through the Amouslek Formation and
Abb. 11. Thin-section of archaecyath bioherm from the Amouslek Formation. Detail shows transverse section of the archaeocyath Neoluculicyathus cf. abadiei (DEBRENNE 1959) the cup of which is almost completely filled with calcimicrobes and is vastly overgrown by Renalcis sp. and Epiphyton sp. Amouslek section, northern rim of western Anti-Atlas. Collection of Institut für Paläontologie, Würzburg University.
the abrupt transition into overlying shales of the Issafen Formation are similar to lithofacies changes through Grand cycles in the southern Canadian Rockies (e.g., AITKEN 1966, 1972). This same shale–carbonate transition and an abrupt change into overlying shales are seen near Timoulaye Izder in the western Anti-Atlas. Dark limestones largely replace the shales in the lower part of the formation in the Tiznit region (CHOUBERT 1963; BOUDDA et al. 1979; BERNEKER & GEYER 1990). However, the lower part of the formation in the southwestern and western parts of the Anti-Atlas are dominantly shale (i.e., the „schistes de Timoulay“ of CHOUBERT 1963). In the northwestern Anti-Atlas and parts of the western and southwestern Anti-Atlas, the succession locally consists of massive limestone with thin intercalated shales („Calcaires massifs de Tiznit“ of CHOUBERT et al. 1975). These limestones consist of a framework of calcareous algae with local archaeocyathan build-ups. Depositional environment The vertical succession of the Amouslek Formation is a macroscale (Grand cycle) variation that seems to include a relative sea-level rise or increase in the rate of sealevel rise. This transition is well illustrated at Timoulaye Izder where shale-dominated facies of the lower Amouslek Formation sharply overlie the peritidal, carbonatedominated Igoudine Formation and are progressively replaced upward by trilobite-hash limestones, archaeocyathan-algal build-ups, ooid limestones, and uppermost, cyanobacterial mats with mud cracks. The transition from the trilobite hash limestones into the archaeocyathan build-ups is marked by a 3–4 m-thick interval of slump-folded strata at this section (W. HELDMAIER, E. LANDING & G. GEYER, unpub. data; Stop 17, Field Trip Guide, this volume). TUCKER (1986b) reconstructed mesoscale, peritidal cycles or alternations and transgressive cycles in the upper, carbonate-dominated part of the Amouslek Formation in the northwestern Anti-Atlas. These alternations include a lower, lagoonal-tidal flat unit overlain by a high energy ooid shoal cap interpreted to have been formed during minor sea-level rises. The amount of siliciclastic material in the Amouslek Formation increases to the east (GEYER 1989a). The decrease in development of the carbonate upper part of the Amouslek suggests an easterly transition from the middle carbonate platform into more inner detrital belt facies. Fossils The Amouslek Formation locally has abundant shelly fossils. HUPÉ (1953) and GEYER (1996) described many trilobite species from the formation. Particularly noteworthy are the index fossils of the Fallotaspididae (Fallot-
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aspis, Choubertella, Daguinaspis) and Neoredlichiidae (e.g., Resserops, Marsaisia). Amouslek archaeocyathans include the most diverse assemblages from Morocco and were dealt with in detail by F. DEBRENNE (DEBRENNE 1958, 1959, 1960a, 1960b, 1961, 1964, 1992a, DEBRENNE & DEBRENNE 1965, 1975, 1995, DEBRENNE et al. 1990). The massive limestones of the formation are often dominated by a framework consisting of calcareous algae and subordinate archaeocyathans. These limestones grade into bioherms composed primarily of archaeocyathans with associated trilobites and echinoderm fragments. Other common fossil groups include brachiopods (GEYER 1994a, 1994b), hyoliths, chancelloriids and other SSFs, and calcimicrobes. Biostratigraphically important sections include Amouslek and Tazemmourt, and archaeocyathans are well developed at Amouslek, Tazemmourt, and at the Jbel Taïssa. Stromatolites and low cyanobacterial mats occur in the upper carbonate half-cycle at many localities and have been discussed from Jbel Kissane (BUGGISCH et al. 1978). Synonyms Série schisto-calcaire (CHOUBERT 1943, CHOUBERT & FAUREMURET 1970); Série des Schistes verts et Calcaires intermédiaires (HUPÉ 1953); Série schisto-calcaire de l’Anti-Atlas (CHOUBERT 1956); Série schisto-calcaire ou Étage d’Amouslek (BOUDDA et al. 1974); Étage Amouslekien (CHOUBERT et al. 1975); Slate-limestone formation (pro parte) (HOLLARD 1985); Slates and sandstones Formation (HOLLARD 1985; = probably a writing mistake); Amouslek Formation (GEYER 1989a, 1990b).
Lemdad Formation Lithology and regional development The Lemdad Formation (GEYER 1989a, 1990b) is a geographically restricted sequence of brownish, red, and
Fig. 12. Top of Lemdad Formation at Lemdad-Le I section. Collaborator W. Heldmaier stands on top of a surface formed by archaeocyath–calcimicrobial limestones (= Ounein B sensu DEBRENNE & D EBRENNE , 1995). Note domal tops of build-ups) overlain by green shales and patches of volcaniclastic material.
greenish marlstones and siltstones with thick limestones and dolostones and volcanic ash or volcanoclastic intercalations. It is described only from the „bloc occidental“ of the „Massif ancient“ (i.e., ‘Adrar n’Dren’ or ‘Promontoire d’Ouzellarh’) of the High Atlas. The lithologic variability of this unit is due to the vertical changes in volcanic content. Its thickness is estimated to be 500 m in the Lemdad syncline. The type section at Oued El Mdad (section Le I; Stop 4, field excursion guide, this volume) consists of a relatively rhythmic alternation of marl- and siltstones with limestones and dolostones. The lower part of the formation is rich in oolitic and oncoidal limestones and dark lime mudstones. The carbonates are replaced upward by finely crystalline dolostones and light colored, somewhat sparitic limestones (BOUDDA et al. 1975; GEYER 1989a). This change is accompanied by a decrease in carbonate bed abundance and thickness. The average thickness of the carbonate-dominated intervals is about 30 to 40 m. Towards the top of the formation, archaeocyathans form conspicuous build-ups at Oued El Mdad, but they do not extend laterally (GEYER et. al. 1995). The fine-grained siliciclastic beds are variable, but are largely composed of unfossiliferous slaty siltstones that are brown, red, or grey in the lower part and increasingly grey or green upward. Thin, often nodular marlstones and less abundant, micaceous or feldspathic quartz arenites are present. Basic to intermediate volcanic intercalations mask the seeming cyclicity of the type section. The volcanoclastics occur as distinct beds within the fine-grained siliciclastics, but are less easily distinguishable within carbonate-dominated intervals. Spilites and keratophyres are known as boulders (BOUDDA 1968) and probably represent clasts reworked from lower levels of the formation.
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The Lemdad Formation of the central High Atlas overlies the Lie de vin Formation, underlies the Issafen Formation, and correlates with the Igoudine and Amouslek Formations. GEYER (1989a) proposed the Lemdad Formation because this post-Lie de vin and pre-Issafen interval could not be subdivided on the available evidence into an Igoudine and Amouslek interval, as in the Anti-Atlas, and because of the importance of volcanic interbeds. It is possible that further work may allow recognition and correlation of the carbonate–shale break at the Igoudine– Amouslek contact in the Anti-Atlas into the Lemdad Formation. This correlation would allow recognition of volcanic-rich facies of the Igoudine and Amouslek Formations and suppression of the Lemdad Formation.
part of the Lemdad (SDZUY 1978), and Dipharus (=Hebediscus), Bondonella, Antatlasia, Berabichia, and many others appear in the middle and upper part (GEYER 1988b, 1990c). Fallotaspis occurs in green slates near the Tizi n’Test (GEYER 1989a). Additional shelly fossils include brachiopods, hyoliths, Hyolithellus, and various phosphatic microfossils. Archaeocyathans form large bioherms in the middle to upper part of the formation, and have been partly described by DEBRENNE et al. (1992) and DEBRENNE & DEBRENNE (1995). „Calcareous algae“ (e.g., Epiphyton, Girvanella, and Renalcis) are frequently intergrown as a framework in the archaeocyathan build-ups. Synonyms
Depositional environment Little investigation of the paleoenvironments of the Lemdad Formation has been done. The formation represents shallow marine depositional environments with varying siliciclastic and volcanic input. Wave activity was moderate, but stronger ambient energies are indicated by occasional coarser beds. Field observations suggest that the formation, at least in its upper part, is a regressive or aggradational mesoscale cycle that is best illustrated by carbonate microfacies.
Série schisto-calcaire p. p. (CHOUBERT 1952b, 1958); Série schisteuse (CHOUBERT 1952b, CHOUBERT 1956); Complexe schisteux (CHOUBERT in HUPÉ 1953); Étage d’Issafene (CHOUBERT 1963, BOUDDA & CHOUBERT 1972, CHOUBERT et al. 1975); Assise schisteuse ou étage d’Issafène (BOUDDA et al. 1974); Étage Issafenien (CHOUBERT et al. 1975); Sous-étage d’Issafene (BOUDDA et al. 1979); Green and purple slate and rubbly limestone formation (HOLLARD 1985); Green and purple slates and Scoriaceous limestones Formation (HOLLARD 1985, probably a writing mistake); Schistes de l’Issafene (SIEGERT 1986); Schistes de l’Issafène (BUGGISCH & SIEGERT 1988); Lemdad Formation (GEYER 1989a, 1990b). Other names which are probably spelling errors are not listed.
Fossils Details on the fossil sequence are known only from the Lemdad syncline, but most of the species are still undescribed. SCHMITT (1979a, 1979b) described the endemic stromatolites Vetella safartiae and Madiganites lemdadensis from the lower part of the formation in the Lemdad syncline. Trilobites are common in some horizons through the unit. Several species of Lemdadella occur in the lower
Tislit Formation Lithology and regional development The Tislit Formation (GEYER 1989a, 1990b) crops out east of the Lemdad, Igoudine, and Amouslek Formations in the central and east-central Anti-Atlas and has its westernmost outcrops in the Tazenakht area. It is a heterolithic unit of alternating, red or purple siltstones and
Fig. 13. Tislit Formation at Id Bouktir, central Anti-Atlas. The section features typical parasequences that commence with thick dolostones with abundant low algal mats in the upper part, overlain by siltstones with ripple marks and cross-bedding, and ending with shales with halite crystal marks and desiccation cracks.
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Fig. 14. Sketch map of the High Atlas and Anti-Atlas regions, Morocco, illustrating tentative facies distribution during deposition of the lower Issafen, uppermost Lemdad, upper Tislit, and Akerouz Formations and tentative paleogeography during their deposition. The line that starts at Agadir indicates a zone of thrust faults, which separate the Anti-Atlas domain from the Hercynian fold belt. Irregular wavy pattern represents non-depositional areas; large stipples = tentative distribution of Tislit Formation; spotted brick signature = tentative distribution of Lemdad Formation; horizontal hachures = Issafen Formation; brick signature = areas with significant amounts of Calcaire scoriacé beds in the Issafen Formation. Based on maps from GEYER (1989a: Fig. 5) and CHOUBERT and on unpublished data from G. GEYER and W. HELDMAIER.
fine-grained quartz arenites with thick intervals of largely dolomitic carbonates. The siltstones are generally argillaceous but range into fine-grained, feldspathic quartz arenites towards the east. Greenish shales that occur in the formation are commonly laminated and become rare to the east. As for other Lower Cambrian units in southern Morocco, siliciclastic content increases toward the basin margin in the east and north (VAN LOOY 1985; GEYER 1989a). However, medium- to coarse-grained sandstones appear even in the westernmost sections of the Tislit Formation in the Tazenakht region. Ripple marks, cross-beds, dessiccation cracks, and large mud clasts occur in the sandstones (BUGGISCH et al. 1978; VAN LOOY 1985). Pseudomorph halite crystals are known from the cover of the El Graara massif (DESTOMBES, unpublished). The lateral changes in lithology accompany an easterly thinning of the formation from a maximum estimated thickness of more 300 m to its feather edge in the eastern Jbel Sarhro or at the Jbel Ougnate (VAN LOOY 1985; BOUDDA et al. 1979).
Sedimentary cycles easily recognizable and particularly well developed south of the Jbel Sarhro (BUGGISCH et al. 1978). Reddish or grey, often calcareous dolostones dominate the carbonate beds of the formation, particularly in the lower part. VAN LOOY (1985) described numerous oolitic and intraclast limestones from the upper part of the formation in the Hmam and Tislit-n-Tamassine synclines. These limestones were correlated with a carbonate interval with low cyanobacterial mats in the Tislitn-Ait-Douchchêne syncline. BUGGISCH et al. (1978) described an interval of black, bituminous limestones, grey dolostones, and green and red marlstones and shales just below the lowest occurrence of trilobites at sections south of Jbel Kissane. These „Calcaires de base“ were interpreted to mark an interval of more normal marine salinity and were regarded as an analog and correlative of the Tiout Member. VAN LOOY (1985) assumed a volcanic origin for the abundant detrital quartz in the calcareous dolostones at the type section. In addition, BUGGISCH et al. (1978) described crystal tuffs from the lower part of the formation
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in the Jbel Sarhro. In general, however, volcanic ash is an insignificant component of the formation. Depositional environment Sedimentary cycles (parasequences) in the Tislit Formation south of the Jbel Sarhro indicate extremely shallow marine depositional environments (BUGGISCH et al. 1978). These alternations include stromatolitic carbonates that are overlain by siltstones with cross-beds and ripples. Dessication cracks point to subaerial exposure. The top of the cycles are composed of shale with halite pseudomorphs formed under hypersaline conditions (SIEGERT 1986). Oolitic limestones with intraclasts in the western part of the formation’s outcrop belt suggest high energy shoal environments that may be transitional laterally into the Igoudine Formation. These environments were replaced to the east by lagoonal facies and supratidal flats. Fossils Skeletal fossils are relatively rare in the Tislit Formation. Trilobites, hyoliths, and calcareous algae are known from a number of calcareous horizons through the formation, but none of them have been formally described. Domal and mat-like stromatolites and thrombolites are abundant in parts of the formation in the Jbel Sarhro (BUGGISCH et al. 1978). Synonyms Série schisto-calcaire, à facies orientale (CHOUBERT 1952b); Calcaires supérieurs [pro parte] (BUGGISCH et al. 1978); Calcaires de base [pro parte] (BUGGISCH et al. 1978, BUGGISCH & FLÜGEL 1988, Série schisto-calcaire [pro parte] (BUGGISCH et al. 1978); Série schisto-calcaire orientale (CHOUBERT et al. 1975, BOUDDA et al. 1979); Série Schisto-calcaire orientale (VAN LOOY 1985); Série des calcaires supérieurs [pro parte] (BUGGISCH & FLÜGEL 1988); Tislit Formation (GEYER 1989a, 1990b).
Issafen Formation Lithology and regional development The Issafen Formation (GEYER, 1989a) consists of grey, yellowish, reddish, and purple slaty shales and siltstones with only a few calcareous and/or quartz arenite beds in the lower part and rubbly or nodular limestones (Calcaire scoriacé) in the upper part. Thin volcanic ashes occur in the unit. Rocks of the formation are widely exposed along the rim of the Anti-Atlas and the southern slope of the High Atlas. The known thickness ranges from 26 m in parts of the Tazenakht region (VAN LOOY 1985) to 100 m between Akka and Tata and up to 180 m at Amouslek. The lower part of the Issafen typically consists of slightly micaceous, finely laminated, light shale and siltstone, often with a slaty cleavage as particularly developed in the southwestern part of the Anti-Atlas. The fresh color of the rock is greenish to light yellow and weathers to very light grey to white, sometimes with light yellow to light brown lines that mark calcareous laminae. The fine-grained siliciclastics are usually planarlaminated, with local cross-bedding and burrow churning. Fossil grainstones and marlstones are locally present, but limestones are a minor component in the sequence. Towards the top of the Issafen Formation, increasingly common siltstones and fine-grained quartz arenites and rare volcanic ashes appear. The Issafen Formation has local archaeocyathan build-ups that may form topographically conspicuous ribs within the soft slaty shales and siltstones. Closely packed calcareous nodules with a shale matrix (i.e., ‘scoriaceous limestones’) are confined in the Issafen Formation to the western Anti-Atlas. These calcaires scoriacé beds form the base and top and thin intervals within the formation. Calcaire scoriacé has been interpreted as diagenetically altered beds of fine-grained ter-
Fig. 15. Amouslek section at the northern flank of western AntiAtlas. The rocks which form the lower and middle slope are the Issafen Formation and show conspicuous units of rubbly limestones (Calcaire scoriacé beds). The top of the hill is formed by resistant sandstones of the lower part of the Tatelt Formation.
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rigenous clastic and subordinate volcanoclastic sediment and calcareous interbeds (DE KONING 1957; HUPÉ 1959). However, a more probable origin lies in diagenetic redistribution of calcareous material within green, purple, and red siliciclastic muds and the growth of nodules within burrow-churned sediment (e.g., LANDING 1995b). The Issafen Formation coarsens into red siltstones and intercalated, fine-grained quartz arenites of the Tislit Formation in the central Anti-Atlas. Restudy of the series of shales with subordinate limestone beds at the type section of the „Serie von Jbel Tichinchine“ (SIEGERT 1986) near Timoulaye Izder in the western Anti-Atlas proves extreme lithologic similarity with the Issafen Formation and indicates that the „Jbel Tichinchine Formation“ is lithologically comparable and correlative with the Issafen Formation (GEYER & LANDING 1995). Fossils The remarkable archaeocyathans of the Issafen Formation have been examined by F. DEBRENNE (DEBRENNE 1964; DEBRENNE & DEBRENNE 1975, 1992a; DEBRENNE et al. 1992) at the famous Amagour and Tamezrhar–Aguerd localities (DEBRENNE & DEBRENNE 1965). The first reported African archaeocyathans were discovered in 1925 in massive bluish limestones of the formation at Sidi Moussa d’Aglou (BOURCART 1927; BOURCART & LE VILLAIN 1928a, 1928b, 1931). In addition, hyoliths, brachiopods, echinoderms, chancellorids, and cyanobacterial mats are known from the formation (BERNEKER & GEYER 1990; GEYER, unpublished data). Synonyms Série schisto-calcaire p. p. (CHOUBERT 1952b, 1958); Série schisteuse (CHOUBERT 1952b, CHOUBERT 1956); Complexe schisteux (CHOUBERT in HUPÉ 1953); Grès noiratres (in part) (HUPÉ 1959); Étage d’Issafene (CHOUBERT 1963, BOUDDA & CHOUBERT 1972, CHOUBERT et al. 1975); Assise schisteuse ou étage d’Issafène (BOUDDA et al. 1974); Étage Issafenien (CHOUBERT et al. 1975); Sous-étage d’Issafene (BOUDDA et al. 1979); Green and purple slate and rubbly limestone formation (HOLLARD 1985); Green and purple slates and Scoriaceous limestones Formation (HOLLARD 1985, probably a spelling error); Schistes de l’Issafene (SIEGERT 1986); Schistes de l’Issafène (BUGGISCH & SIEGERT 1988); Issafen Formation (GEYER 1989a, 1990b); Grès noirâtres [pro parte] (HUPÉ 1959); Serie von Jbel Tichinchine (SIEGERT 1986); Series [Série] of Jebel Tichinchine, Série d’Jebel Tichinchine (BUGGISCH & SIEGERT 1988); additional names which are probably spelling mistakes are not listed.
Tazlaft Formation Lithology and regional development The Tazlaft Formation (GEYER 1989a, 1990b) is a monotonous series of pinkish, reddish or purple, locally green-
grey, fine- to medium-grained, feldspathic sandstones and arkose with rare, red argillaceous siltstones. The typical arkoses and sandstones are medium- to massively bedded and show conspicuous trough cross-beds. The Tazlaft Formation was originally introduced for the outcrops in the east-central and eastern Anti-Atlas (BUGGISCH et al. 1978; BOUDDA et al. 1979), and was noted to thin towards the east and south. The Tazlaft was earlier interpreted (GEYER 1989a) to represent the eastern facies of the supposedly terminal Lower Cambrian sandstones traditionally called the „grès terminaux,“ while the socalled Asrir Formation (abandoned, see LANDING et al., 2006) was interpreted as the western facies of the „grès terminaux.“ Reevaluation of the „grès terminaux“ now shows that it is composed of two successive, lithologically distinct sandstones separated by a depositional sequence boundary. As now understood, the Tazlaft Formation comprises the lower „grès terminaux,“ and is a unit which can be identified across nearly the entire Anti-Atlas. This lower „type Asrir“ in the western Anti-Atlas is simply a distal marine, burrow-churned, open-shelf facies of the Tazlaft. The Tazlaft Formation is now understod to underlie the Tatelt Formation, which had been assigned traditionally to the top of the Asrir Formation, and thus is simply the lower part of the „grès terminaux“ (LANDING et al. 2006). The Tazlaft Formation is usually conformable on the Issafen Formation, but locally lies on Précambrien III (BOUDDA et al. 1974). The type section section of the Tazlaft Formation near Tazlaft is 115 m thick, and the Tinifift locality has the greatest known thickness, 160 m (SIEGERT 1986). In other east-central Anti-Atlas localities, the Tazlaft Formation is less than 20 m thick. The Tazlaft Formation is absent in the central High Atlas, apparently as a result of block faulting and uplift in the early Middle Cambrian, and the volcaniclastic sandstones of the Tatelt Formation unconformably overlie the Issafen Formation (LANDING et al. 2006). Tazlaft formation sandstones are frequently trough cross-bedded, 0.2-3.0 m-thick units that form lenses up to 15 m-thick (SIEGERT 1986). These lenses frequently show basal lag deposits and represent channel-fill and accretionary point bar deposits. Dessication cracks have been described only from the lower part of the formation, where the characteristic cross-beds are not developed (SIEGERT 1986; BUGGISCH & SIEGERT 1988). Unidirectional paleocurrents toward the west to west–southwest in the southern Jbel Sarhro and toward the west to westnorthwest in the El Graara massif are indicated by the cross-beds (SIEGERT 1986; HELDMAIER, 1997). The Tazlaft Formation in the western Anti-Atlas mainly consists of alternating shale and sandstone intervals that are several tens of meters in thickness. The red,
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Fig. 16. Well bedded, cross-stratified quartz arenites of the Tazlaft Formation at Tizi n’Telgane (between Jbel Sarhro and Jbel Ougnate).
olive-green, and green-grey shales have large amounts of fine-grained quartz sand and may be highly micaceous. The associated sandstones are green-grey or light green, quartz-rich or feldspathic, and often highly micaceous. Lithic arenites („greywackes“) and arkoses higher in the section often have large pyroclastic or, rarely, ferruginous patches. Depositional environment The Tazlaft Formation marks an important change in deposition in southern Morocco, as it records an influx of sandstones above the mudstone-dominated Issafen Formation. SIEGERT (1986) and BUGGISCH & SIEGERT (1988) proposed a depositional setting that was controlled by two deltaic complexes. These were regarded as braided river systems, even though the depositional structures of the relatively fine-grained sandstones of the Tazlaft Formation are better regarded as the product of a meandering flow regime. A macrotidal sandfield facies with ebb flow-dominated structures has ben identified at Ourika Wawrmast (LANDING et al., 2006). This evidence for marine deposition means that interpretations of the Tazlaft as a riverine deposit are misleading. The influx of sands marked by the Issafen Formation
took place in the earliest Middle Cambrian. LANDING et al. (2006) equated this sand influx to regional offlap to eustatic regression during the Hawke Bay regression. A supposedly transitional facies between the Tazlaft and the earlier identified Asrir Formation has been regarded as marking a change into a tide-dominated delta system (BUGGISCH & SIEGERT 1988). Now, this part of the former Asrir Formation is regarded a part of the Tazlaft Formation distinguished simply by a change in facies to sublittoral, finer-grained, more intensely burrow-churned sandstone (LANDING et al., 2006). Siliciclastic mudstone clasts, cross-beds, and ripple marks are typical for the Tazlaft Formation in the western Anti-Atlas (SIEGERT 1986; BERNEKER & GEYER 1990). Deposition there records an increase in coarser siliciclastic debris over the Issafen Formation. In most parts of the western Anti-Atlas, coarsening- and fining-up sequences are intrepreted as evidence for progradational events of a delta complex (see Aguerd Formation). However, it is unlikely that the area was part of a wave-influenced, nearshore system. Parts of the formation with occasional cross-beds and quartz arenites with Skolithos-like traces may represent littoral sands (BERNEKER & GEYER 1990). A few pyritic sandstones with abundant Skolithoslike, vertical burrows extend over wide areas in the south-
27
western Anti-Atlas (HUPÉ 1959; DESTOMBES 1985; BERNEKER & GEYER 1990). Planolites-like traces are locally abundant. The thickest sections of the Tazlaft Formation in the region of Alnif have low energy calcarenites and siltstones. SIEGERT (1986) proposed that these sediments indicate limnic conditions (‘Lake of Alnif’), although low energy, inner shelf conditions seem more appropriate.
(p. p.) (CHOUBERT 1958); Série gréso-calcaire (p. p.) (MOUSSU 1959); Étage d’Asrir (p. p.) (CHOUBERT 1963, BOUDDA & CHOU-BERT 1972, CHOUBERT et al. 1975); Grès terminaux ou étage d’Asrhir (p. p.) (BOUDDA et al. 1974); Asririen (p. p.) (CHOUBERT et al. 1975); Niveau d’Asrir (p. p.) (BOUDDA et al. 1979); Terminal slates, sandstones, and tuffs formation (p. p.) (HOLLARD 1985); Terminal slates sandstones and tuffites Formation (p. p.) (HOLLARD 1985, probably a spelling mistake).
Aguerd Formation Fossils Locally, the formation yields vertical trace fossils (Diplocraterion, Skolithos, Teichichus). Body fossils are not reported. Synonyms Grès terminaux [p. p.] (CHOUBERT 1952b, CHOUBERT in HUPÉ 1953, CHOUBERT & FAURE-MURET 1970, BOUDDA et al. 1979); Grès roses à influences continentales (CHOUBERT et al. 1975, BOUDDA et al. 1979); Grès rouges (p. p.) (CHOUBERT et al. 1975, BOUDDA et al. 1979); Grès rouge (SIEGERT 1986); Grès rouges (BUGGISCH & SIEGERT 1988); Asrir Formation (p. p.) (GEYER 1989a, 1990b, GEYER & LANDING 1995); Tazlaft Formation (GEYER 1989a, 1990b); Série gréseuse et volcanique
Lithology and depositional environment The thin (10-20 m) Aguerd Formation (SIEGERT 1986) includes cross-bedded and quartzose, coarse-grained arenites and conglomerates. The formation crops out on the southern margin of the western Anti-Atlas from the Tata area west to the area of El Aïoun du Draa and east into the Jbel Ougnate area, with the type locality at the mouth of the Aguerd valley some distance from the eponymous village. Planar bedding changes into asymmetrical ripples and trough cross-beds upward through this massively bedded unit. The cross beds indicate northerly transport.
Fig. 17. Sketch map of the High Atlas and Anti-Atlas regions, Morocco, illustrating tentative facies distribution during deposition of the Tazlaft, Aguerd, and Akerouz formations and a tentative paleogeography during their deposition. The line that starts at Agadir indicates a bundle of thrust faults which separate the Anti-Atlas domain from the Hercynian fold belt. Irregular wavy pattern = non-depositional areas; large stipples = tentative distribution of upper Akerouz Formation; medium-sized stipples = tentative distribution of the eastern facies of the Tazlaft Formation; small regular stipples = areas with coarser-grained Tazlaft Formation; minute stipples areas with finer-grained Tazlaft Formation; irregularly stippled = distribution of Aguerd Formation during deposition of the Tatelt Formation. Arrows indicate directions of sediment transport. Based on maps from GEYER (1989a: Fig. 6), SIEGERT (1986), HELDMAIER (1997), and CHOUBERT, and on unpublished data from G. GEYER.
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Burrow mottled layers and Skolithos occur only in the lowest part of the formation. The recognition and utility of the Aguerd Formation in mapping is problematical because the formation was based primarily on granulometric study (SIEGERT 1986). Indeed, characters that might be appropriate for confident definition remain undefined. Fossils Trace fossils, almost exclusively Skolithos tubes („pipe rock“), from the lower part of the formation are the only known biological remains. Synonyms Grès verts marins (CHOUBERT et al. 1975, BOUDDA et al. 1979); Konglomerat[e] von Aguerd (SIEGERT 1986); Conglomerate[s] of Aguerd, Conglomerate d’Aguerd (BUGGISCH & SIEGERT 1988); Aguerd Formation (GEYER 1989a, 1990b, GEYER & LANDING 1995).
Akerouz Formation Lithology and regional development The Akerouz Formation (SIEGERT 1986) is known only from the eastern Anti-Atlas at localities 60 km WNW of Erfoud. The formation was first assigned to the „Grès rouges“ (BOUDDA et al. 1979) and includes 70 m of white and grey quartz arenites and purple siltstones and mudstones at the type locality. According to SIEGERT (1986), the lower 35 m are a heterolithic alternation of grey to yellowish sandstone, microconglomerate, and reddish shale units that are up to several meters in thickness. The upper part is characterized by grey and whitish sandstones. Large-scale cross-beds, scour surfaces, ripples, thin laminations, and mud pebbles within the sandstones are reported by SIEGERT (1986). The Akerouz Formation rests on Precambrian or Cambrian(?) sandstones and conglomerates. SIEGERT (1986) believed that the Akerouz was a condensed equivalent of the entire Lower Cambrian elsewhere in Morocco, with the calcarenite bed at its base correlative with the deepening event at the Amouslek–Issafen formational contact. All of these hypotheses require testing by biostratigraphic study and lithostratigraphic correlation. Depositional environment SIEGERT (1986) interpreted this unit as the product of fluvial deposition under the influence of northerly to northwesterly flow. BUGGISCH & SIEGERT (1988) reported caliche in the Akerouz Formation and proposed that subaerial conditions are recorded. Fossils No fossils are known from the Akerouz Formation.
Synonyms Grès rouges [pro parte] (BOUDDA et al. 1979); Serie von Akerouz (Géorgien continental) (SIEGERT 1986); Série von Akerouz (BUGGISCH & SIEGERT 1988); Série d’Akerouz (BUGGISCH & SIEGERT 1988); Akerouz Formation (GEYER 1989a, 1990b).
Tatelt Formation Lithology and regional development The Tatelt Formation (LANDING et al., 2006) is a thin (to 31 m), heterolithic unit. The rock color in lower half is green, light gray, and red siltstone with trilobite-bearing, hematitic limestone or calcareous sandstone beds. The upper half generally consists of feldspathic quartz arenite–arkose with hummocky cross-beds. This formation is an erosionally resistant, cliff-forming succession which represents most of the upper part of the traditional „grès terminaux,“ or Asrir Formation (abandoned, LANDING et al., 2006). The Tatelt Formation has a broad outcrop along the margin of the Anti-Atlas and also occurs in the „Bloc occidental“ of the „Massif ancien“ in the High Atlas. Marker beds and parasequence boundaries allow tentative correlations of these thin Tatelt sections south and west into the thick sections near the Souss Basin axis. Lithofacies and bounding surfaces define three informal members of the Tatelt at its type section that can be regionally correlated and suggest the unit is composed of four parasequences. Differences in sandstone coarseness and the amount of volcanic material result in variable lithology and color of the Tatelt Formation. Volcanic activity seems to have reached a maximum during deposition of this unit; the centers of volcanism are believed to lie in the western High Atlas region or somewhat farther north and northwest in the early Middle Cambrian (earlier thought to be late Early Cambrian by CHOUBERT & FAURE-MURET 1956; SCHAER 1964; BOUDDA et al. 1979; BUGGISCH & SIEGERT 1988). Multiple eruptions have been considered responsible for sedimentary alternations with a lower unit of dark green, slaty ashes (SIEGERT 1986) but this hypothesis awaits confirmation by field data. Depositional environments The Tatelt Formation typically shows readily distinguishable, erosive lower and upper contacts, and comprises a distinct lower Middle Cambrian depositional sequence at the top of the „grès terminaux.“ As the formation records marine onlap, the „grès terminaux“ can no longer be considered the record of a simple regression. Rather, the Tatelt Formation records transgression with sand influx and accumulation during a dramatic interval of basin reorganization.
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Epeirogenic activity that included uplift and allowed deep erosion and removal of the Tazlaft Formation from a tilted block in the Lemdad area on the south slopes of the High Atlas, beveling of the Tazlaft Formation in the central Anti-Atlas, and redefinition of the axis of the Souss Basin in the western Anti-Atlas all preceded Talelt Formation deposition and led to development of a type 1 sequence boundary at its base. The marine shelf deposition shows aspects of variable facies. Small ripples in lower member 2 show the persistence of wave-dominated, shelf environments. The upward transition of sedimentary structures in the sandstones of member 3 (i.e., burrow-churned — thin-bedded — HCS — structureless red sandstone with nodules) is regarded as a shoaling-up sequence from open marine shelf through shoreface deposits that ends with peritidal sandstones. Fossils The abrupt change in trilobite faunas marked by the appearance of the lowest Middle Cambrian (lower Agdzian Stage) Hupeolenus Zone is indicative of the Tatelt Formation and indicates a distinct difference to the underlying Tazlaft Formation. Numerous trilobites are found in a calcareous, tuffaceous arkose at Amouslek. Several other localities yield limited faunas with trilobites (HUPÉ 1953, 1959, 1960; GEYER 1990a, 1990b). In addition to trilobites, helcionelloid mollusks, hyoliths, and brachiopods have been found. Synonyms Grès terminaux p. p. (CHOUBERT 1952b, CHOUBERT in HUPÉ 1953, CHOUBERT & FAURE-MURET 1970); Série gréseuse et volcanique p. p. (CHOUBERT 1958); Série gréso-calcaire p. p. (MOUSSU 1959); Étage d’Asrir p. p. (CHOUBERT 1963, BOUDDA & CHOUBERT 1972, CHOUBERT et al. 1975); Grès terminaux ou étage d’Asrhir p. p. (BOUDDA et al. 1974); Asririen p. p. (CHOUBERT et al. 1975); Niveau d’Asrir p. p. (BOUDDA et al. 1979); Terminal slates, sandstones, and tuffs formation p. p. (HOLLARD 1985); Terminal slates sandstones and tuffites Formation p. p. (HOLLARD 1985, probably a spelling mistake); Série volcano detritique d’Jbel Issendalen [sic] (SIEGERT 1986); „Série volcano-détritique of Issendalen“ (BUGGISCH & SIEGERT 1988); Asrir Formation p.p. (GEYER 1989a, 1990b).
Feijas internes Group The Feijas internes Group (DESTOMBES 1985), based on the „Schistes de Feijas intérnes“ (CHOUBERT 1942), is an informal lithostratigraphic term (i.e., a name that is not based on a geographic or mapped cultural feature) for lower–middle Middle Cambrian rocks that underlie an elongate lowland (i.e., feijas) in the Akka–Tata type area. DESTOMBES (1985) distinguished a lower, siltstone and fine sandstone-dominated Jbel Wawrmast Formation and
Fig. 18. Lemdad syncline, upper part of section Le I (see Stop 4, field excursion guide, this volume). Lower half of slope formed by Issafen Formation with rhythmic intercalations of arenitic limestone beds, upper slope by arenitic and frequently volcanodetrital deposits of the Tatelt Formation.
an upper, sandstone-dominated Jbel Afraou Formation in the central and eastern Anti-At-las. Three laterally correlative formations are established west of Akka: the shaly Tamanart Formation („Lower shale formation of the inner Feijas“), the overlying, resistant sandstones of the Goulimine Formation, and the upper shaly Akka Formation („Upper shale formation of the inner Feijas“) (GEYER 1989a). Recent investigations suggest that an interfingering of these facies occurs and a clear regional differentiation of the lithostratigraphic units may not be possible (HELDMAIER 1997). More importantly, the „Tamanart Formation“ is regarded simply as a deep-water facies of the Jbel Wawrmast Formation that was deposited near the axis of the Souss Basin. Thus, „Tamanart Formation“ has been abandoned and is now regarded as a junior synonym of the Jbel Wawrmast Formation (LANDING et al. 2006). The base of the Jbel Wawrmast and Tamanart Formations (and base of the Feijas internes Group) is discordant and is a depositional sequence boundary (LANDING et al. 2006). The transgressive character of the base of the group is best demonstrated in the central High Atlas and northern Jbel Sarhro, where the group rests nonconformably on the Proterozoic basement. That area
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may represent either: 1) a persistant morphologic high following the late Panafrican collision of a plate in the north with the West African craton that was first inundated by the Feijas internes Group, or 2) a block uplifted in the late Early Cambrian by a change in the strike-slip tectonic regime on this part of margin of the West African craton’s margin. Jbel Wawrmast Formation Lithology and regional development The typical Jbel Wawrmast Formation (DESTOMBES 1985) largely consists of yellowish-weathering, fine-grained sandstones and green and blue-green siltstones and mudstones. The lower part of the unit is characterized by thin, lithologically variable, bedded to nodular limestones interbedded with thin siliciclastic mudstones and siltstones. The part of the formation with the multiple, usually bioclastic limestones is termed the Brèche à Micmacca Member (described below), and crops out from the western High Atlas to the eastern and west-central Anti-Atlas. The greatest thickness of the Jbel Wawrmast Formation is more than 300 m in the Tagragra syncline southeast of Tazenakht (VAN LOOY 1985). This thickness
decreases to approximately 100 m and less in the Jbel Ougnate (DESTOMBES 1985). A monotonous sequence of green, grey-green, and bluish shales and fine-grained sandstones with a few intercalated quartzitic beds has earlier been distinguished as the Tamanart Formation (GEYER 1989a). These coeval strata are now regarded as a lateral equivalent of the typical Jbel Wawrmast Formation and constitute a junior synonym. Arkoses are locally developed as units up to several meters in thickness. Authigenic glauconite is frequently present. The darker colors of the Tamanart facies by comparison to the siliciclastic rocks of the Jbel Wawrmast Formation is either explained by the stronger Hercynian deformation of the western Anti-Atlas (BUGGISCH 1988) or by a higher primary quartz content of the sediments. A few calcareous horizons that probably represent the Brèche à Micmacca Member are present in the lower part of the Tamanart facies, but the lithosequences of several sections (Amouslek, Ifrane d’Anti-Atlas, Akka Tsem) indicate that the lowest equivalents of that member are well above the base of the formation so that the diachronism of the lower boundary is obvious. The thickness ranges from 700(?) m in the southwestern-most
Fig. 19. Sketch map of the High Atlas and Anti-Atlas regions, Morocco, illustrating distribution of the Feijas intérnes Group. The line that starts at Agadir indicates a zone of thrust faults which separate the Anti-Atlas domain from the Hercynian fold belt. Irregular pattern represents non-depositional areas; hatchured area indicates Anti-Atlas embayment that represents regions without Lower Cambrian deposits that were flooded during the early Middle Cambrian transgression. Irregularly hachured = areas with significant amounts of volcanic detritus. Isopach lines indicate probable thicknesses of the Feijas internes Group (in meters). Based on maps from DESTOMBES (1985), GEYER (1989a: Fig. 7), and HELDMAIER (1997).
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Anti-Atlas to more than 400 m at Aguerd and to 250 m near Akka (DESTOMBES 1985). Synonyms Lower shale formation of the inner Feijas (DESTOMBES 1985); Jbel Wawrmast Formation (DESTOMBES 1985); Tamanart Formation (GEYER 1989a, 1990b).
Brèche à Micmacca Member The Brèche à Micmacca Member is the lower part of the Jbel Wawrmast Formation, and locally, the „Tamanart Formation.“ It is composed of frequently trilobite-, echinoderm-, and brachiopod-rich, bedded and nodular limestones that alternate with shale or sandstone intervals. The limestone beds often cap color cycles that extend upward from yellowish green to bluish and finally red. Thin (up to 5 cm) volcanic ashes (frequently K-bentonites in the High Atlas and Jbel Sarhro) occur in the member, and pillowed flows lie at the base of the member locally in the Jbel Sarhro (LANDING et al. 2006). Reworked clasts of the underlying siliciclastics, glauconite grains, submarine corrosion surfaces, and stromatolitic laminae frequently occur in the thin, frequently amalgamated limestone (up to 40 cm) beds. A number of these reddish to greenish shelly limestones can be interpreted to mark erosive intervals that developed during offlap-onlap events. The member crops out from the eastern and eastern Adrar n’Dren massif of the High Atlas to the western Jbel Sarhro region (Agdz and Skoura areas) and the El Graara massif. The recognition of this informally named member was based on particularly fossiliferous limestones exposed near Ourika Wawrmast on the main road from Ouarzazate to Agdz, 100 km to the southeast of Tizi n’Tichka in the north-central Anti-Atlas (BONDON & NELTNER 1933). The Brèche à Micmacca Member commences with a few
Fig. 20. Oued Boutergui OB1 section, Jbel Kissane area. The gully exposes a typical suite of fine-grained sandstones and fossiliferous, limonitic limestone beds of the Brèche à Micmacca Member. Note Jbel Wawrmast in the distance (see Fig. 21).
meters of purple or grey tuffitic and somewhat arenaceous slates in this „type“ section. The overlying interval has rhythmically intercalated ferruginous limestones (Fig. 20) that are up to several decimeters in thickness and have abundant shelly fossils in a superb preservation. The thickness and lithology of these beds varies greatly. The name „Brèche à Micmacca“ was originally applied to a lithology that bears a superficial resemblence to a breccia formed of intraclasts, fossil fragments, and local calcite crystals. Calcareous, tuffaceous, dark shale and conglomeratic or oolitic iron ore clasts occur, as well. These limestones normally include well-washed, fossil grain- to wackestones composed largely of trilobite sclerites and brachiopod valves. Alternations of grey to green shales and limestones follow higher in the section. These carbonates are rather thin and vary in lithology and fossil content. A second, thick, condensed carbonate interval lies roughly 20 m higher in the formation. Impure limestones, up to 0.5 m-thick, consisting partly of trilobite coquinas, are separated by green shales, with purple to red shale tending to underlie the limestone directly. The number of limestones and thickness of the member vary greatly. The „Niveau d’Ouriken-n’Ourmast“ (CHOUBERT et al. 1975; BOUDDA et al. 1979) was described by BUGGISCH et al. (1978) from south of the Jbel Kissane. There, a thick carbonate unit with a trilobite coquina is intercalated in the upper part of the member. The thickness of the member depends on the distance to the basin margin and on the development of carbonate beds in the upper part of the member. For these reasons, thicknesses of the member vary greatly, with VAN LOOY (1985) reporting 76 m in the Ighels syncline and between 20 and 40 m present in the Lemdad syncline (see Stops 5 and 7, Field excursion guide, this volume).
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Fig. 21. Jbel Wawrmast, central Anti-Atlas. Rocks composing the lower and middle slope are the Jbel Wawrmast Formation. Areas with color cycles in the lower part and in the foothills in the foreground are the Brèche à Micmacca Member. The top of Jbel Wawrmast is formed by the Jbel Afraou Formation. The type sections of the Jbel Wawrmast Formation and its Brèche à Micmacca Member are in a transverse gully tothe left of the picture.
Tarhoucht Member (new) The Tarhoucht Member is proposed here as the second, and upper, member of the Jbel Wawrmast Formation. The relatively thin (91 m) type section is located in the eastern Anti-Atlas on the northern slope of Bou Tiouit, about 1 km southwest of the village of Tarhoucht at Lambert coordinates 536/485. GEYER et al. (1995, p. 112–114, fig. 25) have detailed the lithology of the type section near the Tarhoucht village. „Tarhoucht“ has not been used previously for a formal lithostratigraphic interval in Morocco. HELDMAIER (1997) defined „Tarhoucht Member“ in the same way as used herein in an unpublished doctoral dissertation. At Tarhoucht, as at most localities in southern Morocco, the Tarhoucht Member overlies the Brèche à Micmacca Member; comprises the majority of the Jbel Wawrmast Formation; is succeeded by the Jbel Afraou Formation, with the base of the Jbel Afraou defined by the lowest bed with HCS; and is a relatively homogeneous interval characterized by dominantly gray to green colored, yellowish-green weathering, very finegrained sandstones, siltstones, and mudstones. Minor nodular and shell hash carbonates; thin volcanic ashes low in the member; and fossiliferous, coarser-grained
sandstones or arkoses high in the member may be present. The base of the Tarhoucht Member at most localities (i.e., central and eastern Anti-Atlas, south slope of the High Atlas) is a distinctive, lithostratigraphically abrupt change that involves the loss of the highest fossiliferous, bedded and nodular carbonates in Morocco (e.g., GEYER et al. 1995). In the latter areas, limestones interbedded with reddish or purplish mudstones comprise the Brèche à Micmacca Member, and are abruptly replaced by gray, greenish, or bluish, fine-grained sandstones, siltstones, or mudstones of the Tarhoucht Member. In coeval, deeper-water successions along the axis of the Souss Basin in the western Anti-Atlas (e.g., Oued Boutergui, Aguerd, Amouslek), the Brèche à Micmacca is replaced by finegrained, distal tempestites and turbidites with alternating green, purple and red colors and scattered calcareous nodules (e.g., GEYER et al. 1995: 65, 67), and the base of the Tamanart is defined by a transition into laminated, green or greenish medium-grained quartz arenites and arkoses. The Brèche à Micmacca–Tarhoucht contact is diachronous in the lower Middle Cambrian across southern Morocco, because the upper Brèche à Micmacca Member ranges locally into the lower Ornament-
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aspis frequens to Kymataspis arenosa Zones (GEYER 1990c; GEYER & LANDING 1995, p. 41). The appearance of laminated, medium-grained sandstones frequently characterizes the uppermost Tarhoucht Member, and thin HCS sandstones at the base of the Jbel Wawrmast Formation define the top of the Tarhoucht Formation in most Moroccan successions. The Tarhoucht–Jbel Afraou contact is markedly diachronous because these HCS sandstones locally appear as low as the upper Cephalopyge frequens Zone and as high as the Kymataspis arenosa Zone (GEYER et al. 1995). These HSC sandstones have not been recognized in the Lemdad region on the southern slope of the High Atlas (e.g., GEYER et al. 1995, p. 91), and the Jbel Formation and Tarhoucht Member facies persists into the higher middle Middle Cambrian. Facies and depositional environments Deposition of the uppermost Lower and lowermost Middle Cambrian rocks under the Jbel Warwmast Formation took place took place in on a sandstone dominated marine shelf (Tazlaft and Tatelt Formations) in the eastern Anti-Atlas (discussed above). The base of the Brèche à Micmacca marks a strong marine incursion and the onset of widespread, shallow, low energy siliciclastic platform conditions that differ greatly from the underlying, higher energy, and higher depositional slope environments of the underlying sandstones. This interval of basin reorganization, which featured the development of local volcanic flows, deposition of a number of K-bentonites, and onlap across the Pan-African orogen in the High Atlas and Jbel Sarhro (CHOUBERT et al. 1975; DESTOMBES 1985; GEYER 1989a; GEYER & LANDING 1995; LANDING et al., 2006), is considered epeirogenic and not eustatic in origin. Marine onlap progressing eastward first led to the deposition of the Tamanart facies. The greenish, relatively homogenous aspect of the lower Tamanart sandstones represents a low energy, siliciclastic shelf or ramp. The arkosic arenites indicate a source area that did not contribute to the deposition of the Jbel Wawrmast facies in the east. Paleocurrent data from the Jbel Taïssa suggest a sediment transport from the east or southeast. Occasional HCS in these sandstones indicates a deeper shelf environment. The marine onlap advanced from west to the east. By the end of Jbel Wawrmast deposition, the Adrar n’Dren massif was covered, and southern Morocco was completely inundated by shallow seas. The green to purple or red alternations of the Brèche à Micmacca Member have a limestone cap with erosive base, and this suggests an interval of mesoscale offlap-onlap events with condensed horizons that are tempestites at the top.
Fossils The Jbel Wawrmast Formation represents most of the classic, trilobite-rich „Schistes à Paradoxides“. Ironically, the name „Brèche à Micmacca“ refers to a genus which is not present in the fauna. Acanthomicmacca is found with several different species in the limestone beds, but it is questionable if BONDON & NELTNER (1933) even applied the name to a species of Acanthomicmacca. In addition to trilobites (HUPÉ 1953; GEYER 1988a, 1988b, 1990a, 1990c, 1994c) and various groups of brachiopods (MERGL 1982, 1988; GEYER & MERGL 1995), echinoderms, hyoliths, chancelloriids (SDZUY 1969), mollusks, and other small shelly fossils are present (GEYER 1986; LANDING et al. 1995), along with low planar and tiny domal stromatolites on the limestones. Archaeocyathans, mentioned to be found in the „Brèche à Micmacca“ of the central Anti-Atlas (BUGGISCH et al. 1978), come from reworked clasts in sandstones of the underlying Tazlaft Formation. Fossils are sparse in the Tamanart mudstone facies. Trilobites from the Cephalopyge and Ornamentaspis frequens zones are known from Amouslek, and taxa most probably of the Ornamentaspis frequens Zone have been recovered at Aguerd. Age The formation is early Middle Cambrian (early Agdzian Age). It rests on the Tazlaft, Tatelt, and Akerouz Formations with erosional contacts and overlies progressively older Early Cambrian units so that it locally rests on the rocks affected by the Pan-African orogeny. Synonyms Dalle à Micmacca, Passée calcaire mince à Micmacca ellipsocephaloides et Protolenus Latouchei (sic) (BONDON & NELTNER 1933); Brèche à Micmacca (B ONDON & N ELTNER 1933, CLARIOND & GUBLER 1937); Calcaires inférieurs (p. p.) (NELTNER 1938); Brèche (Dalle) à Micmacca (CHOUBERT in HUPÉ 1953); Niveau d’Ourmast (CHOUBERT 1963); Niveau à « brèches à Micmacca » (CHOUBERT & F AURE -M URET 1970); Niveau d’Ouriken-n’Ouarmast (CHOUBERT & FAURE-M URET 1970, BOUDDA et al. 1979); Niveau d’Ouriken n’Ouarmast (= Ourmastien) (C HOUBERT et al. 1975); Niveau schisteux d’Ouriken-n’Ouarmast (BOUDDA et al. 1974); Ourmastien (p. p.) (BOUDDA et al. 1979); Grès d’Ourmast (p. p.) (SIEGERT 1986); Brèche à Micmacca Member (GEYER 1989a, 1990b), Jbel Wawrmast Formation (GEYER 1989a, 1990b).
Jbel Afraou Formation Lithology and regional development The Jbel Afraou Formation (DESTOMBES 1985) largely consists of greenish to yellowish, fine-grained sandstones with intercalated coarser-grained quartz arenites. The content of coarse detrital material increases from the northwest to the southeast (DESTOMBES 1985; HELDMAIER
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1997), and the siliciclastic material has been interpreted to have arrived at the southern rim of the Souss Basin from a delta complex (CHOUBERT 1952b). The formation includes a fossiliferous, argillaceous, fine-grained sandstone at Jorf Dahl in the northeastern Jbel Ougnate (DESTOMBES 1985, and unpubl. data). The formation is identified from the south-central Anti-Atlas to the Jbel Ougnate region in the eastern Anti-Atlas. It has a maximum thickness of 400 m at the margin of the El Graara massive and thicknesses of 200 m and 100 m, respectively, east of Agdz and in the Jbel Ougnate area (DESTOMBES 1985, HELDMAIER 1997, and unpublished data). The quartz arenites exhibit a variety of sedimentary structures that include cross-bedding, hummocky crossstratification, and channels. The 10-30 m-thick sandstone-dominated intervals of the formation show lensing of the depositional units. Shale clasts are present in these sandstones. Calcareous layers, including arenaceous trilobite coquinas and rare nodules, are present. Depositional environment The base of the Jbel Afraou Formation can be defined by the lowest occurrence of hummocky cross-stratified sandstones (HELDMAIER 1997). These sandstone beds mark a transition from a low energy shelf into a higher energy, possibly lower shoreface lithofacies assemblage. In accordance with fairly depauperate faunas, three to four packages of hummocky cross-stratified sandstones characterize sections through the Jbel Afraou Formation. Local fluvial environments may be present (HELDMAIER 1997). These data indicate that the Jbel Afraou Formation marks a regressive or aggradational sequence. Fossils The well-washed quartz arenites may have Skolithos burrows (i.e., „Tigillites“ sensu DESTOMBES 1985), and body fossils are generally absent from the coarser siliciclastics. However, the quartz arenites at the Jbel Arhouri section in the northern Jbel Sarhro, yield such trilobites as Kingaspis and „Paradoxides“, and bradoriids. Fine-grained sandstones and arenaceous shales and siltstones, which often resemble the Jbel Wawrmast Formation, locally contain well-preserved faunas that permit precise biostratigraphic assignment. A fauna with such trilobites as Eccaparadoxides, Ctenocephalus, and Badulesia tenera and the orthid brachiopod Brahimorthis antiqua (HAVLÍCEK 1971) has been found at Hassi Brahim in the west-central Anti-Atlas (DESTOMBES 1985; G. GEYER, unpubl. data; see Stop 14, Field Excursion Guide, this volume). An older trilobite fauna with Condylopyge, Ornamentaspis, and Eccaparadoxides is known from
fine-grained sandstones at Jorf Dahl (GEYER 1988a, 1990c, unpubl. data). Synonyms Grès de Zagora p. p. (CHOUBERT 1952b); Sandstone formation of the Jbel Afraou (DESTOMBES 1985); Jbel Afraou Formation (GEYER 1989a, 1990b).
Goulimine Formation Lithology and regional development The Goulimine Formation (DESTOMBES 1985) consists of resistant, cliff-forming, coarse-grained quartz arenites with local horizons with Skolithos tubes (i.e., „Tigillites“ in the earlier literature). It crops out from El Aïoun du Dra to the Akka area in the western Anti-Atlas. However, exact correlation of the formation is not yet clear, and the beds identified as Goulimine Formation on the Bou Izakarn map sheet are most probably not coeval with those assigned to the formation on the Foum-el-Hassane map sheet. Lithologic and stratigraphic comparison suggests that the formation is partly coeval with the Jbel Afraou Formation. Depositional environment Depsitional environments of the Goulimine Formation are poorly studied. The report of Skolithos in the formation (DESTOMBES 1985) and our observations of low-angle trough cross-sets, current ripples, and small shale clasts in the locally burrow-mottled, glauconite sandstones of the Goulimine Formation suggest marine, tidally-influenced deposition. Synonymy Goulimine formation (DESTOMBES 1985), Goulimine Formation (GEYER 1989a, 1990b).
Akka Formation Lithology and regional development The Akka Formation (GEYER 1989a) is a relatively monotonous sequence of grey, bluish, and grey-green shales and quartz arenites that resembles the Jbel Wawrmast and Tamanart Formations. Fine-grained sediments dominate the base of the formation and grade up into sandstone complexes in the upper part of the formation. The thickness ranges from about 200 m in the extreme southwestern Anti-Atlas to 325 m near Akka (DESTOMBES 1985). Depositional environments Depositional environments of the Akka Formation remain poorly studied. The formation is a coarsening-up, aggradational or progradational, inner shelf package.
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Synonyms Upper shale formation of the inner Feijas (DESTOMBES 1985); Akka Formation (GEYER 1989a, 1990b).
in Morocco (DESTOMBES & FEIST 1987; GEYER et al., 2002). Tremadocian deposits unconformably overlie the Jbel Lmgaysmat Formation. Synonyms
Tabanite Group The subdivisions of the sandstone-dominated Tabanite Group (DESTOMBES 1985) are poorly studied, and their contacts are not yet formally defined in this relatively monotonous succession of sandstones and finer-grained siliciclastics. The group is best exposed along the southern rim of the Anti-Atlas, where it forms a conspicuous ridge flanked by the less resistant Feijas internes Group and the Ordovician Feijas externes Group. It is named for the conspicuous ridge known as the Jbel Tabanite between Akka and Tata. Greyish green sandstones and grey or local red quartz arenites characterize the group. The upper part is usually green or light grey quartzite. The group ranges from a maximum thickness of 380 m in the northwestern AntiAtlas, where the group is not divisible into formations recognizable elsewhere, and thins from 350 m at Agadir Tissint to 55 m at N’Koub (Jbel Sarhro) and measures only 5 m in the Adrar-Zougar borehole (DESTOMBES 1985). These thickness differences are due to partial to nearly complete erosion before the Early Ordovician transgression. Four formations may be distinguished in most of the outcrop area (particularly the central and eastern AntAtlas) and represent alternations from sandstone-dominated to shale and sandstone intervals. Only two formations are distinguishable further southwest in the Anti-Atlas. Most of the Tabanite Group is Middle Cambrian in age. The Jbel Lmgaysmat Formation at the top of the succession has the only Upper Cambrian trilobite faunas
Fig. 22. Tabanite Group in the Alougoum section, south-central Anti-Atlas. The slope consists of upper Rich Khlifa (resistant sandstones at the base), Bailiella (more or less argillaceous sandstones with calcareous beds), and lower Azlag Formations.
Grès à Conocoryphe et lingules; Grès du Tabanit; Conocoryphe and lingulid sandstones (DESTOMBES 1985); Tabanite sandstone group (DESTOMBES 1985); Tabanite Group (GEYER 1989a, 1990b, 1990c).
Rich Khlifa Formation Lithology and regional development The Rich Khlifa Formation (DESTOMBES 1985) con-sists dominantly of resistant, fine-grained quartz arenites capped by an often dark-colored bed of fine-grained conglomerate. The lower part of the formation includes generally more argillaceous arenites. The formation can be traced west of the type locality at the Rich Khlifa mountain (south of the El Graara block) probably into the south-western Anti-Atlas and east to the the western Jbel Ougnate. It ranges from 15 to 75 m in thickness, but differences in thickness are primarily caused by subsequent erosion. The conglomerate at the top of the formation is radioactive due to reworked grains of sphene, zircon, and allanite (DESTOMBES 1985). Depositional environment Depositional environments of the Rich Khlifa Forma-tion remain poorly studied. The unit apparently represents shallow marine, peritidal environments, with possible deltaic or riverine influences. Fossils Fossils (including trace fossils) are unknown from the Rich Khlifa Formation.
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Synonyms Sandstone formation of Rich Khlifa, Rich Khlifa sandstones (DESTOMBES 1985), Rich Khlifa Formation (GEYER 1990c).
Bailiella Formation Lithology and regional development The informally named Bailiella Formation (DES-TOMBES 1985) is a non-resistant unit of green siltstones with a fine-grained argillaceous sandstone interval that forms a conspicuous band. Outcrops of the Bailiella Formation form relatively steep slopes (Fig. 22) or are often covered by alluvial deposits. The Bailiella Formation consists of lower green shales that grade upward into fine-grained, micaceous sandstones with a significant component of shell fragments and mud flakes. Numerous calcareous layers that include marly beds with shell hash („lumachelle“) or arenaceous, calcareous nodules are present and create the banded aspect of the outcrops. A particularly common faunal element of the calcareous is the name-giving trilobite Bailiella levyi. The formation is known throughout the central AntiAtlas, including the Jbel Ougnate massif. Its thicknesses range from 30 to 70 m (DESTOMBES 1985, and unpubl. data), but major differences in thickness are probably a result of subsequent erosion. The sandstones coarsen westward, and the formation is not recognizable west of Hassi Brahim (west of Tata; Stop 14, Field Excursion Guide, this volume). In the Lemdad area of the High Atlas, the formation is probably represented by a lithologic sequence with calcareous layers and nodules with Bailiella levyi and other trilobites, but its thickness is not yet determined (see Stop 7, Field Excursion Guide, this volume).
Depositional environment Depositional environments of the formation are poorly known. Lower energy environments than those of the underlying Rich Khlifa Formation are represented by the Bailiella Formation although progressive wave activity is recorded towards the top of the formation. The fossil hash beds probably represent sorting by storm-wave activity, and the marly layers represent either early diagenetic carbonate dissolution and redistribution or methanogenic nodules. Fossils The numerous intercalated calcareous nodules have a low diversity shelly fauna with trilobites (e.g., Bailiella levyi, Planolimbus spp., and „Paradoxides“ spp.), inarticulate and articulate brachiopods (lingulids, acrotretids, orthids), and echinoderm debris (Cincta, eocrinoids). Particularly beds at the top of the formation record a moderately diverse ichnofossil assemblage (Fig. 23). Synonyms Argillo-arenaceous formation with Bailiella (DESTOMBES 1985), Bailiella Formation (GEYER 1990c).
Azlag Formation Lithology and regional development The Azlag Formation (DESTOMBES 1985) consists of massive sandstones with an upward trend from green, micaceous, and chloritic sandstones to siliceous quartz arenites with herring-bone cross-stratification and Skolithos tubes (termed „pipe rock“ in the older literature) at the top. Trough cross-bedded sandstones often cap the laterally extensive Skolithos sand-stone unit at the top of the formation (especially well developed at the Bou
Fig. 23. Upper part of Bailiella Formation. Upper surface shows unidirectional ripple marks, three different species of simple, horizontal trace fossils and the vertical trace fossil Skolithos. Central Tagragra Syncline, central Anti-Atlas.
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Rbia section, southern El Graara massif). The formation extends from the type locality at the Azlag Pass northwest of Zagora into the south-western Anti-Atlas and east to the central Jbel Sarhro. However, determination of the upper and lower contacts of the formation is difficult through this outcrop belt. In sections where the formation has not been erosionally truncated, its thickness ranges from 50 to 70 m (DESTOMBES 1985). Depositional environment Deposition of this inadequately studied formation apparently took place in shallow, tidally influenced environments, including shoreface. The lithofacies trend from chloritic, micaceous sandstones into massive quartzites at the top tracks a regressive or aggradational trend. Fossils In addition to Skolithos at the top of the formation (Fig. 24) and Planolites-type traces, only rare, lingulid brachiopod fragments have been reported. Synonyms Sandstone formation of Azlag (DESTOMBES 1985); Azlag Formation (GEYER 1990c).
This generally recessive unit consists of soft, dark green shale, argillaceous siltstones, and fine-grained sandstones, often with abundant small mud flake clasts. Thin calcareous layers and calcareous nodules yield fossil fragments. The base of the formation is a conglomeratic or cross-bedded sandstone, which is often strongly limonitized. Depositional environment A shallow marine environment for this incompletely studied unit is indicated by its sedimentary features and the presence of a normal marine fauna. Fossils Thin shell hash layers that probably represent storm wave accumulations have yielded most of the faunas. These early Late Cambrian assemblages (DESTOMBES & FEIST 1987) include a paucispecific fossil assemblage with the trilobites Olentella and Seletella (DESTOMBES & FEIST 1987) and the brachiopods Saccogonum saccatum (HAVLÍÈEK 1971) and Billingsella (MERGL 1983; MERGL et al. 1998). A recently discovered moderately diverse trilobite and brachiopod fauna from the Tagragra Syncline promises improved possibilities for correlation into other regions (GEYER et al. 2002; Fig. 25)
Jbel Lmgaysmat Formation Lithology and regional development The Jbel Lmgaysmat Formation (DESTOMBES 1985) seems to be limited to the area between Hassi Brahim, Foum Zguid, and Tazenakht in the central Anti-Atlas. This is probably a result of erosion in the later Late Cambrian. Its thickness is usually around 35 m, and is truncated at the top.
Fig. 24. Azlag Formation, upper part showing thick-bedded siliceous quartz-arenites with low cross-beds (upper half of the photo) and quartz arenites which are densely penetrated by vertical Skolithos tubes („pipe rock,“ lower part, touched by hammer). Section north of Alougoum, El Graara massif region, central AntiAtlas.
Synonymy Argillo-arenaceous formation of Jebel Lmgaysmat (DESTOM-BES in DESTOMBES et al. 1985), Jbel Lmgaysmat Formation (GEYER 1990a, 1990c).
„Lower Tilemsoun Formation“ Two lithostratigraphic units have been proposed in the upper part of the present Middle Cambrian in the south-
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Fig. 25. Jbel Lmgaysmat Formation. Fossiliferous, bioturbated, limonitic limestone bed in the foreground, overlain by siltstones and fine-grained sandstones. Central Tagragra Syncline (newly discovered locality of GEYER et al. 2002).
western Anti-Atlas by DESTOMBES (1985). The „Lower Tilemsoun Formation“ („Lower Tilemsoun sandstone formation“, DESTOMBES 1985) and „Upper Tilemsoun Formation“ are resistant sandstone-dominated units that form a conspicuous topography in the Tarfaya Province. The „Lower Tilemsoun“ and the „Upper Tilemsoun Formations“ are separated by a thick (250–400 m), recessive interval of green, pyritic, fine-grained, quartz arenite with rare calcareous nodules and shale that DESTOMBES (1985) did not formally name. The „Lower Tilemsoun Formation“ has a thickness of 30 to 60 m in the Tilemsoun area. It consists of lightcolored, massive, fine-grained quartz arenites with rare calcareous nodules and inter-calated shale layers that form a topographically distinct ridge in the type area. Coarse-grained sandstones at the top of the „Lower Tilemsoun Formation“have Skolithos (i.e., „Tigillites“ of DESTOMBES 1985).
The „Lower Tilemsoun Formation“ is overlain by an unnamed recessive interval with Bailiella (DESTOMBES 1985). Available evidence suggests that the „Lower Tilemsoun Formation“ is at least partly correlative with the Rich Khlifa Formation farther to the northeast in the Anti-Atlas. „Upper Tilemsoun Formation“ The „Upper Tilemsoun Formation“ consists of 20 to 40 m of ridge-forming, light colored, massive quartz arenites and has about the same lithology as the „Lower Tilemsoun Formation“. The „Upper Tilemsoun Formation“ is underlain by a thick, recessive interval with Bailiella and is the apparent correlative of and probable junior synonym of the lithologically similar Azlag Formation).
Moroccan Precambrian-Cambrian boundary and problems in stable isotope correlation of the Lower Cambrian The definition and location of the Precambrian–Cambrian boundary in Moroccan sections remain problematical. The boundary lies within or perhaps at the base of a thick interval of sparsely fossiliferous, platform carbonates and associated siliciclastics. This interval includes the Adoudou Formation and lower part of the Lie de vin Formation. A number of reports have been published over the last decade, and diverse opinions about the location of the boundary have been proposed. This uncertainty has existed because of 1) erroneous interpretations of the limited biostratigraphic data, 2) problematical techniques in the geochronology of Moroccan volcanic rocks, 3) the absence of a reliable
age for the lowest Cambrian until relatively recently, and 4) seemingly simplistic correlations based on non-conventional correlation techniques (i.e., magneto- and chemostratigraphic logs) of sequences on separate Cambrian continents without adequate biostratigraphic controls or the recognition of major stratigraphic breaks in local sections. Discussions on the Moroccan Precambrian– Cambrian boundary have featured arguments on: 1) differing radiometric dates, mainly on the underlying basement, in the 1970s (CHOUBERT et al. 1965, 1973, 1975, 1979; CHARLOT et al. 1969, 1970; CHOUBERT & FAURE-MURET 1972, 1977; JUERY et al. 1974, 1975; CHARLOT 1976; LANCELOT 1977; DUCROT & LANCELOT 1977; LEBLANC &
39
LANCELOT 1980; CLAUER et al. 1982; MIFDAL & PEUCAT 1985; BUGGISCH & FLÜGEL 1988) and 2) beliefs about the biostratigraphic utility of stromatolites and thrombolites about 1980 (SCHMITT 1978, 1979a, 1979b; BERTRAND-SARFATI 1981; AITKEN & NARBONNE 1989). These arguments are not reviewed below; however, it might be said that the boundary proposals and correlations in these reports reflected a geochronologic and biostratigraphic understanding of the Proterozoic–Cambrian boundary that has been superceded only after the Cambrian lower boundary Global Stratotype Section and Point has been defined in the Fortune Head section in southeastern Newfoundland (LANDING 1994). Many Moroccan geochronologic dates remain controversial at present and are of little value in discussion of the Proterozoic–Cambrian transition. We have undertaken a collaboration with S. A. Bowring that includes high resolution U/Pb dating of volcanic zircons from the Ouarzazate and the Taroudant Groups in the Anti-Atlas and western High Atlas ranges. Although the samples are not yet completely investigated, data on ashes from the upper Lie de vin Formation and from the Issafen Formation (LANDING et al. 1998) contributed to establish correlations between several Cambrian continents (see BOWRING et al. 1992; ISACHSEN et al. 1994). Definition of the Moroccan Precambrian–Cambrian boundary is complicated in comparison with other parts of the world because subtrilobitic, Small Shelly Fossil (SSF) assemblages and the biostratigraphically important trace fossils that locally appear below the lowest SSFs (e.g., LANDING 1994) are not known from Morocco. As discussed above, the restricted marine, frequently hypersaline environments of the Adoudou and Lie de vin Formations largely precluded colonization of these nearshore facies by terminal Ediacaran and earliest Cambrian trace-producing and skeletalized metazoans. Although a few authors have cited a personal communication from the late B. DAILY, who claimed to have found SSFs near the base of the Lie de vin Formation on a field trip to Morocco in 1976 (BOUDDA et al. 1979; KIRSCHVINK 1992), their presence has never been confirmed, and participants on that trip doubt his observations (M. SCHMITT and K. SDZUY, personal communication to G. GEYER). The Ediacaran–Cambrian boundary is bracketed by U–Pb radiometric age determination of rhyolitic lavas from the upper Ouarzazate Group. As mentioned above, such felsic lavas and ash flow tuffs throughout the AntiAtlas were all dated between 565 and 560 Ma (MIFDAL & PEUCAT 1985; WALSH et al. 2002; MALOOF 2004). A recent date from the Minount rhyodacite (the second youngest lava flow northwest of Timjich) had a concordia age of 562.89 ± 0.49 Ma (MALOOF et al., 2005). No reliable radiometric age determination is known from the lower and middle Adoudou Formation. However,
MALOOF et al. (2005) reported a concordia age of 525.38 ± 0.46 Ma for an ash in the upper Adoudou Formation from the Oued Sdas section. This date is very significant for three reasons. 1) The first is that this age indicates that at least this part of the upper Adoudou Formation is significantly younger than the Cambrian lower boundary. 2) The dated ash lies in an interval of a strongly positive d13C excursion correlated with the base of the Siberian Tommotian Stage, and allows for the first strong basis for intercontinental correlation correlation between the West Gondwanan margin and the Siberian Platform. 3) This Moroccan date is younger than the 530.7 ± 0.9 Ma age determined by ISACHSEN et al. (1994) on an ash from southern New Brunswick on the Avalonian microcontinent and biostratigraphically correlated into the lowest Cambrian Manykaian (=Nemakit-Daldynian) Stage. Interregional correlation further indicates that this ash lies somewhat higher than diverse SSF faunas in eastern Newfoundland and southern New Brunswick (LANDING et al. 1989; LANDING 2004). Thus, these Avalonian Small Shelly Fossil assemblages are sub-Tommotian- and Manykaian-equivalent, and represent an assemblage that is Earth’s oldest known skeletalized fauna. Limited biostratigraphic control on the Lie de vin Formation is based on calcareous algae and trace fossils. The biostratigraphic significance of these reports from the Lie de vin Formation at the Tiout section (LATHAM 1990; LATHAM & RIDING 1990) are problematical. As seen on this trip and documented by MONNINGER (1979), the Lie de vin Formation at Tiout is dominated by dolomitized, thrombolitic and stromatolitic carbonates with minor siltstone intercalations and represents a restricted marine facies. Primary environmental, lithofacies, and diagenetic factors can be presumed to have limited the occurrences and recovery of the calcareous algae and the few Cambrian trace fossils (such as Diplocraterion; LATHAM & RIDING 1990) from the section, while the ranges of the few known fossils from Tiout have relatively extended ranges in the subtrilobitic and trilobite-bearing Lower Cambrian. An asserted equivalency of these Tiout faunas with the upper Tommotian of Siberia (e.g., KIRSCHVINK et al. 1991) is speculative. We have come to regard the SHRIMP-based, U–Pb zircon age of 521 ± 7 Ma at Tiout (COMPSTON et al. 1992) as of relatively little significance in Proterozoic–Cambrian boundary correlations. However, we have also collected an ash in the Tiout section at nearly the same level as the one that COMPSTON et al. (1990, 1992) reported on. This Tiout sample indicated a similar IDTIMS age on volcanic zircons of 522.4 ± 2.0 Ma (LANDING et al. 1998). Very significantly, this date also overlaps a ca. 519 ± 1 Ma date that we have determined on trilobite-bearing rocks from the uppermost Lower Cambrian (i.e., lower Branchian Series and Botoman-equivalent) Callavia broeggeri Zone in Avalonian
40
South Wales (LANDING et al. 1998). No precise interregional correlation can be drawn with the lowest trilobites and archaeocyathans in Morocco. The lowest occurrence of these forms in the Tiout section seems to be governed by a change to somewhat more normal marine facies. These occurrences probably do not represent the first evolutionary occurrence of these forms. It should be pointed out that the unfavorable lithologies below the Tiout Member probably did not permit normal preservation of shelly fossils at Tiout. Undeterminable shelly fragments found in the lower unnamed member of the Igoudine Formation at Tiout include trilobite remains. Time-consuming preparation of the lowest noted trilobite horizon from the Tiout section (SDZUY 1978) has not been done in other sections, and earlier trilobites may prove to be present. By contrast with sequences on other warm water Cambrian continents with carbonate platform sequences, the lowest known trilobites at Tiout appear slightly below, not above as in Siberia, the lowest archaeocyathans. SDZUY & GEYER (1988) argued that the Tiout trilobite assemblage appears to represent particularly old clades and may predate the oldest trilobites on the Siberian Platform. The assumption that the Eofallotaspis Zone trilobites from Tiout are late Atdabanian-equivalent (e.g., KIRSCHVINK et al. 1991; DEBRENNE et al. 1992) remains debatable. It should be noted that the conclusion that Tiout trilobites are particularly old disagrees with magneto- and chemostratigraphic correlations claimed by KIRSCHVINK et al. (1991). These authors equated the lowest Tiout trilobites with the Siberian, middle Atdabanian Pagetiellus anabarus Zone; a correlation that strongly contrasts with unequivocal correlations of younger Moroccan faunas of the Sectigena Zone with part of the Botoman Stage of Siberia (GEYER & PALMER 1995; GEYER 1998, 2005a). Unfortunately, our reconnaissance work shows that SSFs cannot be used to resolve the contradictory claims of the interregional correlation of these trilobites. A number of calcareous algae and cyanobacteria that have been used for claims of Cambrian age occur in the carbonates of the Adoudou, Lie de vin, Igoudine, and Amouslek Formations, but their local stratigraphic ranges and utility in correlation remain limited. Tarthinia, Renalcis, and Kordephyton (see LATHAM & RIDING 1990) are reported from the Lie de vin Formation. Kundatia composita KORDE 1973 was described from the older Adoudou Formation of the central Anti-Atlas (BUGGISCH & FLÜGEL 1988). However, this identification is certainly erroneous because the poorly illustrated type material from Siberia has turned out to represent juvenile archaeocyathans (A. YU. ZHURAVLEV, pers. comm. to G. G.). Nevertheless, an understanding of the relatively limited biostratigraphic potential of the calcimicrobes is now coming to
be appreciated (see MANKIEWICZ 1992). However, the use of Kordephyton to establish an upper Tommotian-equivalency of the upper Lie de vin by KIRSCHVINK et al. (1991, Fig. 3) illustrates a less-than-rigorous attempt at biostratigraphy; the genus is known from single horizons in Siberia and Morocco, and these occurrences cannot provide a direct tie line on which Siberian and Moroccan carbon isotope excursions are hung. In addition, a number of such calcimicrobes are known from strata that are now assigned to the uppermost Proterozoic, rather than to the Lower Cambrian. Similar problems confound the utility of acritarchs, and the report of „Vendian“ acritarchs from the Lie de vin Formation (BOUKHARI et al. 1979; CHOUBERT et al. 1979) is now determined to be incorrect, as the U/Pb date and chemostratigraphy on the underlying upper Adoudou Formation conclusively demonstrates a middle Lower Cambrian (lower Tommotian) correlation into Siberia (discussed above). A further complication in a biostratigraphic definition of the Proterozoic–Cambrian boundary in southern Morocco derives from the recent ratification of a Precambrian–Cambrian boundary global stratotype at the base of the Trichophycus pedum Zone at Fortune Head, southeastern Newfoundland. This definition of the global stratotype means that: 1) the lowest Cambrian includes strata well below the Siberian Tommotian Stage and has the Manykaian Stage (i.e., Nemakit-Daldynian) at its base, and 2) the Tommotian is threfore not lowest Lower Cambrian but considerably younger than the base of the Cambrian (LANDING 1994, 1995a). Available reports on Morocco have assumed that the Siberian Tommotian provided a „standard“ for global correlation of the basal Cambrian and focused on attempted interregional correlations of Moroccan strata into this part of the Siberian Lower Cambrian. This approach has influenced carbon and oxygen isotope-based correlations, as well as magnetostratigraphic investigations. The Tiout and Oued Sdas sections in the western Anti-Atlas have been important in these correlations. More-or-less detailed carbon isotope signatures of the Adoudou and lower Lie de vin Formation were first published by TUCKER (1986a, 1989), MAGARITZ et al. (1991), and KIRSCHVINK et al. (1991). These data were used in attempted correlations with the Yudoma and Pestrotsvet Formations in Siberia (MAGARITZ et al. 1991) and into the South China Platform (BRASIER 1991). In addition, RENNER (1994) reported preliminary results. Recently, the carbon and oxygen isotope data were revised and supplemented by numerous additional data from 15 sections through the Lower Cambrian formations across the Anti-Atlas (MALOOF et al. 2005). These investigations established a new standard curve for the Early Cambrian of Morocco. MALOOF et al. (2005) illustrated that the d13C values of the Tabia Member of the Adoudou
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Formation are difficult to correlate between the sections, but generally vary between -4‰ and 0. The values decrease in the lower Tifnout Member of the Adoudou Formation to a minimum of -6‰ and increase upward to a maximum of +7‰ in the upper Tifnout Member. In the uppermost Tifnout Member, the values decrease again to -3‰ and remain between -3.5 and 0.5‰ for most of the Lie de vin Formation. However, a peak value of 3.5‰ is attained in the uppermost Lie de vin and lower Igoudine formations. The d 13C values presented by MALOOF et al. (2005) appear to indicate relatively stable values at about -0.5‰ for the remainder of the Tata Group. Such carbon isotope signatures of the Adoudou, Lie de vin, and Igoudine Formations at Tiout lead to relatively robust correlations. MALOOF et al. (2005) interpreted the -6‰ nadir at the immediate base of the Tifnout Member as indicating the position of the Ediacaran– Cambrian boundary. Taking into account the radiometric data from the upper Tifnout Member, this would indicate that the deposition of this member took place over a period of more than 15 Ma years which, at first glance seems extremely difficult to explain, but can be accommodated by exceptionally slow accumulation rates of sedimentary rock. TUCKER (1986a, 1989) indicated two strong d 13C excursions at the top of the Adoudou Formation (ca. +7‰ vs PDB) and in the lower half of the Igoudine Formation (ca. +2.5‰ vs PDB; TUCKER 1986a), both confirmed by MALOOF et al. (2005). These excursions have been sug-gested to indicate massive increases in biotic production. However, the data permit differing Lower Cambrian correlations (TUCKER 1986a, 1992). LATHAM & RIDING (1990), KIRSCHVINK et al. (1991), MAGARITZ et al. (1991), and RENNER (1994) reaffirmed these carbon isotope signatures and provided additional data. RENNER (1994) reported two shifts to +4.9 and +5.8‰ vs PDB (recorded as +6‰ by MAGARITZ et al. 1991) in the Adoudou Formation at Oued Sdas. The conspicuous d13C maximum below the top of the Adoudou Formation at Tiout (ca. +7‰ in LATHAM & RIDING 1990: Fig. 1, and MALOOF et al., 2005, but +4.2‰ in MAGARITZ et al. 1991, and +4.6‰ in RENNER 1994) was correlated by BRASIER (1991) with the so-called Dahai d13C maximum that appears in the uppermost Zhongyicun Member and characterizes the thin Dahai Member in the middle Meishucunian Stage of the South China Platform. MALOOF et al. (2005) interpreted this maximum in the upper Tifnout Member as coeval with the Manykaian (Nemakit-Daldynian)–Tommotian boundary in Siberia, which is consistent with dates on a clast from a conglomerate below the lower Tommotian in Siberia (dated at 534.6 ± 0.5 Ma by BOWRING et al. 1993). A second carbon excursion higher at Tiout and in the lower part of the Lie de vin Formation (-2‰ fide MAGARITZ et al. 1991, but -4‰
according to RENNER 1994) was correlated by BRASIER (1991) with the Badaowan d13C minimum in the upper Meishucunian of South China (i.e., correlated into the Badaowan Member of the lower Qiongzhusi Formation). SSF-based correlation of the Badaowan minimum in China is somewhat problematical because the dark grey, dolomitic siltstones of this member are sparsely fossiliferous; the Badaowan Member overlies the thin Dahai Member (less than 2.0 m) with diverse Tommotian-aspect micromollusks and is unconformably overlain by the Botomanequivalent Yuan’shan Member of the Qiongzhushi Formation (see QIAN & BENGTSON 1989; LANDING 1994). A considerable hiatus is generally claimed to be present at the phosphate lag between the Dahai and Badaowan Members (e.g., BRASIER 1991), and this would not seem to favor direct correlation of the carbon isotope curve of the Yangtze Platform with the Anti-Atlas. However, the major stratigraphic break within the middle Meishucunian has been reinterpreted as lying somewhat lower at the abrupt faunal break within the upper Zhongyicun Member between the first and second Chinese SSF assemblages (LANDING 1994, figs. 1, 2B, i.e., major unconformity between Meishucunian A and B). This means that the Dahai maximum and Badaowan minimum simply mark local stages in a depositional sequence with the Dahai maximum corresponding to Tommotian-equivalent onlap and the Badaowan minimum marking an immediately subsequent maximum highstand with a phosphate lag at its base. Of course, this latter interpretation means that the Dahai maximum and Badaowan minimum should not be globally correlatable because they record shallow and deeper environments under a stratified water mass with dysaerobic environments at depth. The assumption that major isotope excursions directly reflect changes in the world ocean, rather than local climatic or epeirogenically driven, offlap-onlap cycles of stratified water masses should be evaluated in attempts at nonconventional correlations. An additional unresolved problem lies in the fact that available carbon isotope data on the Moroccan Lower Cambrian first came from two adjacent sections, and regional variations have not been evaluated until recently. RENNER (1994) investigated the carbon isotope curves of the Tiout and Oued Sdas sections, along with sections from the western and central Anti-Atlas. Significantly, the positive d13C excursions of the Tiout and Oued Sdas sections could not be detected elsewhere in Morocco. Even lithologically comparable parts of the sequence that are possible correlatives had distinctly different isotope values. RENNER (1994) suggested that: 1) diagenetic overprint may have distorted a primary isotopic signal assumed to be present or 2) similar lithofacies within the Adoudou and Lie de vin Formations are strongly diachronous. A strong diagenetic alteration of the Adoudou
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and Lie de vin Forma-tions has been shown at Tiout by MONNINGER (1979). In addition, the basin architecture and biostratigraphic data from overlying units (discussed below) suggests that diachroneity may be involved. However, it should be noted that no detailed microstratigraphy and section-by-section correlation of potential marker beds has ever been attempted in the Moroccan Lower Cambrian for the relevant strata, and assertions of significant diachroneity in the lowest Cambrian are questionable. MALOOF et al. (2005) reported high frequency oscillations in the d13C equated with a 106 year variability and recorded by often extremely rapid isotopic shifts, which remain largely unexplained. The most conspicuous of these variations is a 6‰ rise in the upper Adoudou Formation at ca. 529 Ma, which has been detected in three sections. Although such a positive excursion could be the result of primary productivity which greatly exceeded respiration or of a short pulse of carbonate weathering that modified the isotopic composition of the ocean, none of these explanations appeas to be a logical explanation for this case. A similarly rapid variation in d13C values is only known from the early Triassic (PAYNE et al 2004), where it marks the period of recovery after the global Permian–Triassic mass extinction. Rapid cycles in the earliest Triassic coincide with pulses during an interval characterized by the extreme rarity of carbonate secreting organisms (PAYNE et al. 2004). The return to stable d13C values in the early Triassic corresponds to the appearance of gastropods. It might thus be hypothesized that the recorded d13C oscillations in the Adoudou Formation may indeed reflect local changes in abundance of early carbonate secreting organisms (MALOOF et al. 2005). Oxygen isotopic data presented by MALOOF et al. (2005) indicate that few trends are discernible for the Taroudant Group due to high-frequency variability and/ or noise with a mean value of -6‰ to -7‰. d13C versus d18O cross plots do not indicate any statistically significant co-variation of these isotopes (MALOOF et al. 2005).
Available evidence for the Lie de vin Formation does not generally seem to permit any degree of certainty in correlations even at the stage-level with Lower Cambrian sections on other Cambrian continents. It is less than questionable whether the carbon isotope variations at Tiout either permit bracketing of the Tommotian–Atdabanian boundary interval or facilitate correlations of the upper Lie de vin and lower Igoudine Formations into the Siberian Atdabanian, as claimed by KIRSCHVINK et al. (1991). Indeed, the biostratigraphic correlation of the upper Tiout section is inadequate for interregional, stage or zone-level correlation: 1) Kordephyton is known only from one horizon in Siberia and Morocco, and this genus’ poorly established range cannot be used for the determination of a Tommotian equivalency. 2) A middle Atdabanian equivalency cannot be claimed for the lower Igoudine Formation on the basis of the endemic trilobites and archaeocyathans; these forms could be older than Atdabanian (discussed above) or younger in the Lower Cambrian. 3) The seemingly progressively positive d13C excursion from the upper Lie de vin into the lower Igoudine Formation shown by KIRSCHVINK et al. (1991, fig. 3) is based on only four analyses through 300 m of section; the long unsampled intervals between samples at Tiout might contain carbon isotope excursions that mimic and match trends elsewhere in the Zhurinsky Mys section. 4) Finally, as noted above, the legitimacy of the Siberia—Morocco—China correlations as proposed by KIRSCHVINK et al. (1991) are brought into question by their report of a 521 ± 4 Ma age on the supposed „upper Tommotian-equivalent,“ upper Lie de vin Formation. However, the ca. 519 Ma age on the upper Branchian (i.e., Botoman equivalent in Siberia) of South Wales (LANDING et al. 1995) suggests that even the upper Lie de vin Formation may prove to correlate into trilobitebearing Lower Cambrian sections on other continents.
Moroccan Cambrian biostratigraphy and chronostratigraphy The trilobite-based, biostratigraphic framework for the Cambrian of the Anti-Atlas was originally proposed by HUPÉ (1952, 1953). His monograph (HUPÉ 1953) can also be considered to mark the beginning of modern Lower Cambrian biostratigraphy. A later, partial revision of the zonal scheme, introduced at the International Geological Congress in Copenhagen (HUPÉ 1960), was partly based on trilobite genera and species that were never formally established. A considerable degree of confusion was introduced into the understanding of the Moroccan Cambrian by a
tradition of equating of litho-, bio-, and chronostratigraphic terminology in this earlier literatre on the Cambrian of Morocco. This led to problems in understanding the true stratigraphic ranges of a number of taxa (see GEYER 1983, 1990a, 1990c). New information, particularly on the lowest trilobite-bearing Lower Cambrian strata (SDZUY 1978) and on the Lower-Middle Cambrian boundary interval (GEYER 1983, 1990b), has required a revision even of HUPÉ’s (1960) later biostratigraphic schemes (see GEYER 1990a). This new biostratigraphic scheme includes biostratigraphic zones from the lowest occurrences of
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trilobites in Morocco into the middle Middle Cambrian. This new zonation for the Moroccan trilobite-bearing Lower Cambrian through Middle Cambrian has been proven to be useful also in Iberia, and provides the basis for a standard for the Cambrian West Gondwana (GEYER & LANDING 2004). Formal zones for the upper Middle Cambrian and the Upper Cambrian of the Jbel Lmgaysmat Formation are not yet established (Fig. 26). The sub-trilobitic Early Cambrian in West Gondwana and Morocco is now termed the Cordubian Series (GEYER & LANDING 2004) and corresponds approximately with the suggested lowermost series of the Cambrian (name momentarily under investigation by the International Subcommission on Cambrian Stratigraphy; BABCOCK et al. 2005, PENG et al., 2006). The trilobite-bearing Lower Cambrian is the Atlasian Series, and includes the Issendalenian and Banian Stages. The entire West Gondwanan Middle Cambrian is assigned to the Celtiberian Series, and includes the Agdzian, Caesaraugustan, and Languedocian Stages (GEYER & LANDING 2004). The occurrence of Paradoxides sensu lato near the base of the Agdzian Stage (and the abandoned Moroccan Tissafinian Stage; GEYER 1990a) suggests a Middle Cambrian age for most of this stage according to the traditional concept of the Lower-Middle Cambrian. The bases for correlation of the Moroccan Lower-Middle Cambrian boundary have been discussed by GEYER & PALMER (1995), GEYER (1998, 2005a).
Issendalenian Stage The Issendalenian Stage (GEYER 1990a) consists of a vertical sequence of four zones (Eofallotaspis, Fallotaspis tazemmourtensis, Choubertella, and Daguinaspis Zones), HUPÉ’s (1953) traditional F. tazemmourtensis through Daguinaspis Zones are largely based on faunas from shales. Fossils, including the well known archaeocyathans (e.g., DEBRENNE & DEBRENNE 1995), from the associated limestones considerably expand the biostratigraphically significant faunal elements. The concept of the stage derives from the frequent occurrence of olenelloid (s.l.) trilobites which represent the fallotaspidid phylogenetic lineage. In addition to the eponymous olenelloid (s.l.) genera, the zones bear early bigotinid trilobites that have not yet been studied in detail, and those of the Eofallotaspis Zone are of particular interest (see SDZUY 1978, 1981). Fallotaspis is not limited to the Fallotaspis tazemmourtensis, Choubertella, and Daguinaspis zones, as earlier indicated by HUPÉ (1953), but is quite common in the higher Antatlasia hollardi Zone and is even found above that zone. This means that the so-called „Fallotaspis zone“ or „Fallotaspis stage“ (FRITZ 1972; REPINA 1986; and subsequent authors) is a vague concept and
Fig. 26. Table of Cambrian chrono- and biostratigraphic units recognized in the Moroccan Atlas regions.
cannot be recognized or correlated outside of Morocco. For example, correlation of the range of Fallotaspis from the Moroccan Cambrian would apparently expand such a „zone“ into the lower Nevadella Zone of the Laurentian Faunal Province and into the Botoman of Siberia (compare PALMER & REPINA 1993, and remarks on the myth of a globally recognizable „Fallotaspis Zone“by GEYER 1991a). In addition, the two index species of the Siberian Fallotaspis Zone have turned out to be representatives of the archaeaspidid genus Repinaella (GEYER 1996). Thus, the Fallotaspis Zone in Siberia does not directly correlate into Fallotaspis-bearing strata elsewhere.
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Banian Stage The Banian Stage (GEYER 1990a) is composed of three vertically successive zones: the Antatlasia hollardi, Antatlasia guttapluviae, and Sectigena Zones. This upper part of the Moroccan Cambrian is characterized faunally by the predominance of ellipsocephaloid trilobites and the frequent occurrence of such eodiscids as Delgadella, Dipharus (=Hebediscus), and Serrodiscus. Holmiid trilobites, previously used as index fossils by HUPÉ (1953), occur sporadically. The Sectigena Zone has trilobites, which correlate with similar faunas in other regions that have been informally termed the „Serrodiscus bellimarginatus-Triangulaspis annio band,“ or „STH band“ (ROBISON et al. 1977; GEYER 1991a, 1991d, 2005a). Archaeocyathans locally form conspicuous bioherms in the Antatlasia hollardi and Antatlasia guttapluviae Zones. However, they disappear within the Sectigena Zone (DEBRENNE et al. 1992, DEBRENNE & DEBRENNE 1995). Archaeocyathan remains in younger strata are reworked.
Agdzian Stage The Agdzian Stage (GEYER & LANDING 2004) is composed in Morocco of four vertically successive zones: the Hupeolenus, Cephalopyge notabilis, Ornamentaspis frequens, and Kymataspis Zones. An earlier chronostratigraphic scheme for Morocco (GEYER et al., 1995) referred this interval to the Tissafinian Stage (consisting of the three lower zones named above) and the Toushamian Stage, the earliest zone of which was the Kymataspis (arenosa) Zone. HUPÉ’s (1952, 1953) „Myopsolenus magnus Zone“ corresponds roughly to the Cephalopyge notabilis Zone, but no equivalent exists in HUPÉ’s zonation for the underlying Hupeolenus Zone. All of the three lower zones are locally very fossiliferous, a condition that has allowed detailed information on these intervals and the ranges of taxa. As a consequence, this stratigraphic interval, which is crucial in understanding the Lower–Middle Cambrian transition in Morocco and interregional correlation of the Lower– Middle Cambrian boundary, is well studied and permits relatively precise interprovincial and even intercontinental correlations. Therefore, this interval is also a key for establishing a global GSSP for the base of a Series 3 (which will replace the traditional Middle Cambrian) elsewhere. Characteristic of the lower part of this stage is the frequent occurrence of ellipsocephalid-protolenoid trilobites. However, the stage does not correspond to the traditional concept of the „Protolenus stage,“ but is probably only equivalent to the middle and upper part of the Avalonian „protolenid-strenuellid interval“ or the „Protolenus Zone“ (GEYER & PALMER 1995).
HUPÉ (1952, 1953) established the „Myopsolenus magnus zone“ (zone VIII) as the top of the Moroccan Lower Cambrian, but this zone clearly belongs to the basal Middle Cambrian in the Scandian and Avalonian sense as indicated by the presence of Paradoxides (Acadoparadoxides). In contrast, the olenelloid trilobite Cambropallas telesto occurs with Acadoparadoxides in the Cephalopyge notabilis Zone (GEYER 1993, GEYER & LANDING 2004). The latter zone can be correlated with a number of intervals outside of Morocco, which are variously placed in the Lower Cambrian (e.g., „Protolenus“ Limestone at Comley, Shropshire) or Middle Cambrian (Acadoparadoxides mureroensis Zone of Spain). Interregional correlation suggests that the lowest Moroccan „Paradoxides“ is definitely older than any other „Paradoxides“ known thus far and that Cambropallas is the youngest known olenelloid trilobite (GEYER 1993, GEYER & LANDING 2004). Kymataspis arenosa is a common trilobite above the Ornamentaspis frequens Zone. Although its range slightly overlaps that of O. frequens and considerably overlaps with that of Badulesia tenera in the upper part of its range, it serves to characterize a biostratigraphic zone. However, the species’ occurrence is lithofacially limited in the southern High Atlas and west-central AntiAtlas by lithofacies. In the upper part of the zone, Parasolenopleura lemdadensis and other species of Parasolenopleura are locally abundant and may characterize a distinct zone. However, imperfect knowledge and regional limitations due to facies control do allow proposal of a formal zone.
Caesaraugustan Stage Only two zones can be identified in Morocco for the earlier period of the Caesaraugustan Stage; these are the Badulesia and Pardailhania Zones. Both zones were assigned to a regional Moroccan Toushamian Stage (GEYER et al., 1995; now abandoned, GEYER & LANDING 2004). Ptychoparioid and paradoxidid trilobites are characteristic of this stage, which is better developed in Iberia, and the name and concept of Caesaraugustan was transferred from the Iberian stratigraphic scheme (see LIÑÁN et al. 1993). The Badulesia Zone is primarily characterized by the occurrence of Badulesia tenera. Other species of Badulesia are extremely rare. In addition, lithofacies changes limit the geographic range of Badulesia tenera in the Atlas regions to the southern High Atlas and westcentral Anti-Atlas, whereas the species is a common faunal element in other regions such as the Moroccan Meseta, Spain, and eastern Massachusetts. The same geographic limitations are seen for Pardailhania, which is almost exclusively represented in the
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Moroccan sections by Pardailhania hispanica. The zone has thus the character of a range-zone. However, the transition from the Badulesia to the Pardailhania Zone is still poorly known. Nevertheless, the occurrence of Pardailhania in layers overlying the Badulesia interval is in accordance with the biostratigraphic zonation of Spain. No body fossils are known from the upper part of the stage in Morocco.
Languedocian Stage and Furongian The terminal Middle Cambrian Languedocian Stage, which is richly fossiliferous with characteristic Solenopleuropsis faunas in Iberia and southern France, has only informal biostratigraphic divisions. The Bailiella Formation features the arrival of marine fossils that are absent above and below. Thus, the characteristic appearance of Bailiella levyi falls within an interval that
was defined by depositional environments. The lower Furongian trilobites of the Jbel Lmgaysmat Formation (DESTOMBES & FEIST 1987) were first described from a single fauna that is not known from other localities. The first samples of that assemblage collected by DESTOMBES included specimens of „Paradoxides“, which are not yet described and probably reflect a part of the uppermost Middle Cambrian unknown elsewhere in Morocco. Recently, a second, more diverse fauna with Upper Cambrian trilobites and brachiopods was discovered in the Tagragra Syncline of the central Anti-Atlas (GEYER et al., 2002; see Fig. 25 herein). This fauna still awaits careful study and has no trilobites in common with the earlier discovered fauna and promises a higher potential for correlation into other regions. In addition to the trilobite faunas, billingsellid brachiopods have proved to bear some potential for correlation (MERGL et al. 1998).
Diachroneity of lithostratigraphic boundaries Relatively synchronous lithofacies changes are traceable between most of the Moroccan Cambrian formations due to deposition in a basin with very low depositional slope and limited influx of coarse siliciclastics. Higher Cambrian units show evidence for diachroneity of formational contacts, and the time equivalency of some of the rock units is questionable. As detailed by GEYER & LANDING (1995), a striking diachroneity appears to exist for the upper boundary of the Lie de vin Formation. The overlying Igoudine Formation at Tiout has yielded the lowest, local occurrence of trilobites and archaeocyathans from the Tiout Member. However, the lowest trilobites in the High Atlas occur near the top of the Lie de vin Formation (BOUDDA & CHOUBERT 1972) or, more probably, in a secondarily redcolored lower part of the Lemdad Formation. The overlying beds of the Lemdad Formation, a succession of shales, limestones, and dolostones, indicate depositional conditions similar to those of the Igoudine Formation in the western Anti-Atlas. However, trilobites from the base of the Amouslek Formation at Tiout appear in the lowermost Lemdad Formation in the Lemdad syncline. These data indicate that either time-equivalent strata of the lower Igoudine Formation probably lie in the Lie de vin Formation in the Lemdad syncline (SDZUY & GEYER 1988) or that the resolution supplied by the trilobites is limited by severe lithofacies controls. The Igoudine–Amouslek formational boundary can be argued to be slightly diachronous even in the western Anti-Atlas due to the different lower levels of the Fallotaspis tazemmourtensis Zone in this lithologic succession. In addition, the Choubertella Zone in the Oued
Boutergui section, western Anti-Atlas, occurs in the basal shale of the Amouslek Formation (BERNEKER & GEYER 1990), whereas Choubertella first occurs well above the base of the Amouslek Formation in other sections in the western Anti-Atlas (HUPÉ 1959). The contact between the Amouslek and Issafen Formations can be shown to be diachronous on the basis of the abundant trilobite faunas. Precise biostratigraphic bracketing of the contact is possible for several sections. The base of the Issafen Formation lies within the Antatlasia hollardi Zone, and probably just below the Antatlasia guttapluviae Zone, at Timghit in the Issafen syncline. In contrast, careful examination of sections along
Fig. 27. Antatlasia hollardi HUPÉ 1953. Latex cast of cranidium. Holotype, Issafen Syncline, Ida ou Drif section. Collection Museum Nationale d’Histoire Naturelle, Paris, MNHN R 50860. Scale bar equals 2 mm.
46
the southern rim of the Anti-Atlas (Tamezrhar, Tadakoust, Aguerd, Tamanart, Adaï, and Asrir) indicates that the base of the Issafen Formation is within the Antatlasia guttapluviae Zone, but ranges from the base of the zone almost to its top. There is a clear trend for a later onset of Issafen Formation de-position in the Anti-Atlas from north to south and, less clearly, from west to east, in other words from the interior to the margin of the basin. This deposition began rather late in the Lemdad syncline near the basin margin. Limited data also suggests that the deposition of the coarse clastics of the Tazlaft Formation (i.e., lower „grés
terminaux“ or lower „Asrir Formation“ in earlier publications) was initiated later towards the center of the basin. A clear diachronism exists for the Brèche à Micmacca Member. The basal limestones of the member range from the Hupeolenus to the Cephalopyge notabilis Zones, with considerably different relative positions within these zones. Similarly, the upper boundary of the Brèche à Micmacca ranges from a position in the low Ornamentaspis frequens Zone to well within strata with Kymataspis arenosa (Stop 14, Field Excursion Guide, this volume), with a higher position in more marginal areas.
Acknowledgments This article is a contribution to the Cambrian Subdivision Project of the International Subcommission on Cambrian Stratigraphy. We are indebted to W. HELDMAIER for the permission to use unpublished results of his investigations, and to E. BERNEKER for help with archaecyathan data.
G. GEYER acknowledges a Heisenberg grant and financial support by the Deutsche Forschungsgemeinschaft during field work in the 1990s. E. LANDING acknowledges field and laboratory assistance from the National Science Foundation (grants EAR 94-15773, EAR 98-05177, and 0116551).
A contribution to the Cambrian Subdivision Project of the International Subcommission on Cambrian Stratigraphy
47
MOROCCO FIELD EXCURSION 2006
Ediacaran–Cambrian depositional environments and stratigraphy of the western Atlas regions GERD GEYER & ED LANDING Addresses of the authors: Dr. G. GEYER, Institut für Paläontologie, Bayerische Julius-Maximilians-Universität, Pleicherwall 1, 97070 Würzburg, Germany, E-mail ; Dr. E. LANDING, New York State Geological Survey, New York State Museum, Empire State Plaza, Albany, N.Y. 12230, U.S.A., E-mail
Introduction This field excursion emphasizes paleoenvironments, litho- and biostratigraphy, sequence boundaries, volcanic ash occurrences, and potential for intercontinental correlation of a number of important Ediacaran and Cambrian sections of the Moroccan Atlas regions. The arrangement of the sections visited during the field excursion from December 2 to 5, 2006, begins on the northern slope of the western Anti-Atlas with the lowest Cambrian at Tiout and the lower part of the trilobite-bearing Cambrian at Tazemmourt. Complete sections through the trilobite-bearing Cambrian in the Lemdad syncline feature the Cambrian in the western High Atlas range. The subsequent drive across the western Anti-Atlas shows aspects of post-Pan-African Proterozoic, siliciclastic facies with archaeocyathan bioherms in the Issafen Syncline, and spectacular slump-folding of the Lie
de vin Formation as well as a siliciclastical-dominated Lower and fossiliferous Middle Cambrian in the Tata region. The final day in the field features the Devonian and the siliciclastic Middle Cambrian along the southern flank of the western Anti-Atlas and carbonate-dominated Lower Cambrian at the western rim of the Anti-Atlas. Information on the stops includes a summary of the lithologic sequence and, if available, a complete list of faunas found at successive levels. The sample numbers with the prefix F refer either to the faunal levels of HUPÉ (1953, 1959) or SDZUY (1978); sample numbers with a comma indicate a precisely measured level above the base of the section or unit; other samples refer to horizons reported by G. GEYER (e.g., 1986, 1988b, 1990c, 1998) or to subsequent collections by the authors and W. HELDMAIER (1997).
Contents December 2. Stop 1. Tiout Section .................................................................................................................................... 50 December 2. Stop 2. Tazemmourt Section ........................................................................................................................ 61 December 3. Lemdad Syncline ........................................................................................................................................... 65 December 3. Stop 3. Cretaceous deposits on the southern flank of the Lemdad Syncline .............................................. 67 December 3. Stop 4. Section Le I ....................................................................................................................................... 68 December 3. Stop 5. Section Le II ..................................................................................................................................... 79 December 3. Stop 6. Section Le IV .................................................................................................................................... 84 December 3. Stop 7. Section Le XI .................................................................................................................................... 86 December 4. Stop 8. Bou Ighir ........................................................................................................................................... 95 December 4. Stop 9. Boutonnière de Ouaoufenrha ........................................................................................................... 95 December 4. Stop 10. Oui-n-Tatayine ............................................................................................................................... 95 December 4. Issafen Syncline ............................................................................................................................................ 96 December 4. Stop 11. Central Issafen Syncline ................................................................................................................. 97 December 4. Stop 12. Southern Issafen Syncline ............................................................................................................ 100 December 4. Stop 13. West of Imitek .............................................................................................................................. 101 December 4. Stop 14. Hassi Brahim ................................................................................................................................ 101 December 5. Stop 15. Icht ................................................................................................................................................ 108 December 5. Stop 16. East of Timoulaye Izder .............................................................................................................. 109 December 5. Stop 17. Timoulaye Izder ........................................................................................................................... 109 December 5. Stop 18. Akhsas Plateau ............................................................................................................................. 111
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Fig. 1. Geological map of the northern rim of the western Anti-Atlas south of Taroudant showing the location of the sections of Tiout (Stop 1) and Tazemmourt (Stop 2). Slightly modified from BOUDDA et al. (1975).
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50
December 2
STOP 1. TIOUT SECTION Setting The name ‘Tiout’ refers to a group of seven villages about 25 km southeast of Taroudant. The village of Igoudine is located at the northern rim of the Anti-Atlas where the roughly north–south trending section ends (Fig. 1). The strata strike more-or-less west–east and dip 25– 35° to the north. The total thickness of the studied sequence, including Adoudou through Amouslek Formations, is about 1,700 m. Significance The Tiout section is certainly the best studied and most referenced section of both the Moroccan Cambrian and, perhaps, the entire West Gondwanan region. The section it important because of this documentation and its relatively complete exposure, as well as its easy accessability. This type sections of the Lie de vin Formation and the Igoudine Formation (including the Tiout Member) provide especially good examples of the very shallow-marine conditions that dominated deposition of these units, and their sedimentologic features will be examined. Complex cyclicity and the conspicuous de-
velopment of cyanobacterial-stromatolitic build-ups will be studied. Although shelly fossils are almost absent below the Tiout Member, the trilobite and archaeocyathan faunas of the Tiout Member are crucial for the Lower Cambrian biostratigraphy of the Atlas region, and their occurrence shows a close relationship to lithofacies. History of research Work on the Tiout section began with studies in the 1970s, which are documented by reports on algal structures (MONNINGER & SCHMITT 1975; SCHMITT & M ONNINGER 1977), archaeocyathans (D EBRENNE & DEBRENNE 1976, 1978, 1995), and trilobites (SDZUY 1978, 1981; LIÑAN & SDZUY 1978; GEYER 1996). The investigations culminated in monographs on sedimentology, geochemistry and depositional environments (MONNINGER 1979) and stromatolites and thrombolites (SCHMITT 1978, 1979a, 1979b). A second phase of research on the Tiout section brought out a number of reports on carbon isotope curves (TUCKER 1986a, 1986b, 1989; KIRSCHVINK et al. 1991; MAGARITZ et al. 1991; RENNER 1994), magnetostratigraphic investigations (KIRSCHVINK et al. 1991), as
Fig. 2. Tiout section. View from top of the Adoudou Formation shows Lie de vin Formation (dark, rhythmically banded outcrops; carbonate-dominated unit known as „Barre de Tata“ is visible in its middle part) and Igoudine Formation (massive beds on nose of mountain). Note Souss plain and southern flank of High Atlas in the background.
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well as calcareous algae and trace fossils (LATHAM 1990; LATHAM & RIDING 1990), and U–Pb zircon geochronology of volcanic ashes (COMPSTON et al. 1990; LANDING et al 1998; MALOOF et al 2005). The isotopic investigations have been supplemented by recently by additional carbon and oxygen isotopic data (MALOOF et al. 2005). Most of the studies targeted the position of the Proterozoic–Cambrian boundary, and the difficulties about this question are discussed in detail by GEYER & LANDING (1995 and this volume). The fundamental difficulties in correlation by conventional biostratigraphic methods because of unfavorable facies conditions suggested that other sections were needed to supply data more pertinent to this question. Indeed, MALOOF et al. (2005) presented correlations into other sections of the Anti-Atlas that put the data from Tiout into broader perspective. However, these revised views of intercontinental correlation actually confirm earlier, but less well-constrainted hypotheses and still allow alternative explanations (see discussion under „Moroccan Precambrian–Cambrian boundary and problems in stable isotope correlation of the Lower Cambrian“ in GEYER & LANDING, this volume). Lithologic sequence (Figs. 3, 4) Adoudou Formation The sequence starts with the Adoudou Formation, which rests unconformably on the so-called Précambrien III (andesites, rhyolites, and ignimbrites). Due to limited exposure, the lower and middle Adoudou Formation has never been studied in detail at Tiout, and the measured section of MONNINGER (1979) includes only the upper 200 m of the Tifnout Member. Crude measurements of the formation suggest a total thickness of roughly 1100 to 1200 m. The formation has been studied in the Oued Sdas section (about 5 km east of the Tiout section). In that section, MAGARITZ et al. (1991) and MALOOF et al. (2005) detected a d13C maximum in the upper part of the Tifnout Member (=Limestone Member of earlier publications), which the authors correlated with a maximum in the uppermost Yudoma Formation on the Aldan River area, and thus with the Nemakit-Daldynian–Tommotian boundary in Siberia. This positive shift (excursion A of LATHAM & RIDING 1990) was confirmed from the Tiout section by RENNER (1994), who indicated a value of 4.6 per mil vs PDB. The following minimum in the uppermost Adoudou Formation at the Tiout and Oued Sdas sections was correlated with a minimum at the base of the Pestrotsvet Formation on the Aldan River, and thus with the base of the Tommotian (MAGARITZ et al. 1991, MALOOF et al. 2005), without a discussion of the obvious gap between the latter formations (see LANDING 1994; GEYER & LANDING 1995, this volume).
T a b i a M e m b e r . The Tabia Member marks the first transgression over the Pan-African basement. In this section, it consists mainly of calcareous rocks that are overlain by shaly siltstones. The poor exposure does not permit precise measurments and investigations. The thickness of the member is probably between 100 and 150 m. T i f n o u t M e m b e r . The Tifnout Member is a monotonous sequence of grey to brownish dolostones and limestones. The lower part of the member is relatively poorly exposed, but the upper part of around 600 m of peritidal dolostones and limestones can be studied. The upper part is dominated by massive dark grey to black dolostones with intercalations of pale to yellowish argillaceous siltstones. The bed thicknesses of the dolostones reach 15 m. The uppermost part, shown in Fig. 3 (X–A), consists generally of pale, bluish-grey finely crystalline limestones with abundant chert nodules. Meterscale, shallowing-upward parasequences are nicely exposed at our first stop. These parasequences generally start with variably colored silstones and marlstones which are overlain by laminated, often wavy-bedded micrites and grainstones, which themselves are capped by cyanobacterial or stromatolitic carbonates. These microbial laminites frequently show such signs of subaerial exposure as tepee structures, desiccation cracks, pseudomorphs after gypsum, etc. (MONNINGER 1979). CHOUBERT (1963) reported Collenia-type stromatolites, but these occurrences have not been confirmed subsequently. Ferrugineous or silicified seams indicate long-term pauses in deposition and mark the parasequence boundaries. Finely interbedded siltstones and dolostones occasionally show slumpfolds or brecciation and indicate instability of the sediment stack. Lie de vin Formation The Lie de vin Formation is a cyclic sequence with a thikness of about 940 m. Although other sections may provide additional stratigraphic and paleontologic data, the Tiout section has been chosen as the type section of the formation because of its thorough investigation (GEYER 1989a). MONNINGER (1979) and SCHMITT (1979a) distinguished three units termed „lower,“ „middle,“ and „upper“. Although these units reflect environmental changes, it has not yet been possible to trace them into other sections with adequate precision. L o w e r m e m b e r (Fig. 3: A–B): The (informal) lower member has a thickness of about 430 m. Its lower part is designated as a coarsely crystalline dolostone facies, although the unit commences with finely crystalline limestone with chert nodules and thus resembles much of the underlying transitional top of the Adoudou Forma-
52
Fig. 3. Lithologic sequence of the Tiout section (upper Adoudou through Amouslek Formations) with major occurrences of stromatolites, trilobites, and archaeocyathans. Modified from SCHMITT (1979: Fig. 44).
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54
Fig. 4. Lie de vin Formation in the Tiout section. Marlstone-siltstone alternation towards the top of a parasequence in the lower member of the Lie de vin Formation is interrupted by multiple, cm-thick, light colored volcanic ash layers. Pen for scale above the center of the photo.
tion (about 40 m). These limestones have the lowest distinct LLH-type stromatolites known from the section and Morocco in general (Fig. 3: bed 1). Bed 2 in Fig. 3 most probably indicates the lowest recognized occur-rence of Tarthinia- and Renalcis-bearing thrombolites, as mentioned by LATHAM & RIDING (1990). The thickness of this subunit is about 130 m. An overlying facies change is documented by a change to dominant coarsely crystalline dolostones that alternate with purple and greenish shales and so-called ‘argillaceous dolosiltites’ (MONNINGER 1979; Fig. 3: above bed 10). This „purple siltstone and coarse dolostone facies“ is about 150 m in thickness. The dolostones frequently contains nests of „jelly calcite“ that has been interpreted as vug fills (MONNINGER 1979). Thrombolites and columnar stromatolites are known from the dolostones (SCHMITT 1978, 1979a, 1979b). A „pale carbonate facies“ located higher in the lower sequence has a thickness of about 110 m. It is dominated by pink to reddish, coarsely crystalline dolostones interbedded with relatively thick units of thin-bedded argillaceous dolosiltites. The content of medium grey, finely crystalline limestones increases up-section. Small tepee-structures locally occur. M i d d l e m e m b e r (Fig. 3: B–C): The base of the middle member (total thickness about 240 m) is marked by massive black limestones (Fig. 3: beds 12–14) that are believed to coincide with the so-called Barre de Tata (CHOUBERT 1963). These beds are the base of a subunit called „dark limestone facies“ of about 110 m in thickness. This subunit consists of dark grey, locally black, occasionally fetid limestones. Interbeds of shales (argillaceous dolosiltite) are moderately frequent. Although LLH-type and thrombolites form a conspicuous volume of the basal limestone beds (Fig. 3: beds 12–14), more and more thrombolites are found upsection. Thrombo-
lites are most abundant at the top, where up to three build-ups may occur successively within one limestone bed (e.g., Fig. 3: bed 17). The unit shows a transition marked by smaller bedthicknesses and lighter colors, and increasing intercalations of the „pale carbonate facies“ (about 130 m thick). The latter is a variable alternation of grey dolomitic and dark limestones, pale coarsely crystalline dolostones, drab, finely crystalline dolostones and pale shales (argillaceous dolosiltites). Build-ups of columnar stromatolites, cumulate stromatolites and LLH-type stromatolites occur in the dolomitic limestones and are more frequent than thrombolites. The columnar stromatolites include the form genera Tungussia and Patomia. Simple trace fossils occur in the shales, and LATHAM & RIDING (1990: Fig. 2d) probably reported the discovery of Diplocraterion isp. most probably from this subunit. In addition, the subunit likely yielded the tuff layer from which zircons were dated by COMPSTON et al. (1990, 1992) at 521 ± 4 Ma (515 ± 21 Ma). It should be emphasized that the note of COMPSTON et al. (1990) report marked the beginning of a reevaluation of Cambrian age dates. Because of the uncertainty in certain SHRIMP dates, we collected an ash in the Tiout section at nearly the same level. This sample provided a similar IDTIMS age on volcanic zircons of 522.4 ± 2.0 Ma (LANDING et al. 1998). These dates bracket a 519 ± 1 Ma date that we determined on the trilobitebearing, uppermost Lower Cambrian (i.e., lower Branchian Series and a Botoman-equivalent) Callavia broeggeri assemblage in Avalonian New Brunswick (LANDING et al. 1998). U p p e r m e m b e r : The base of the upper member (total thickness about 260 m; Fig. 3: C–D) is again indicated by dark massive limestones. This lithology characterizes a 145 m-thick subunit termed the ‘shale and dark
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Fig. 5. Stromatolite build-up in the upper Lie de vin Formation of the Tiout section. Note transition from low-algal mats in the lower part to domal constructions made up by columnar stromatolites in the upper part of the build-up. Maximum height of build-up ca. 1.3 m.
limestone facies’ (MONNINGER 1979). Nevertheless, the lithologies of this subunit vary over a large scale that starts with dark, finely crystalline limestones that alternate with medium grey dolomitic limestone and drab coarsely crystalline dolostones. The interbeds consist of drab to light purple shales (argillaceous dolosiltites) and olive-green, slaty, calcareous shales. Microbial build-ups reach a maximum within this subunit: Almost every carbonate bed contains stromatolite or thrombolite structures (Fig. 3: beds 30–59). Thrombolites from this subunit bear the calcareous algae Kordephyton, Tarthinia and Renalcis (LATHAM & RIDING 1990). The columnar stromatolites include well preserved specimens of the form genera Nimbophyton, Tungussia, Tifounkeia, and the endemic Anti-Atlas form species Tioutella rupta and Igoudinia hemispherica (SCHMITT 1979a, 1979b). The disappearance of olive shale marks a transition into the „pale dolostone facies“ (about 115 m) with coarsely crystalline dolostones at the base that grade up into drab, finely crystalline dolostones and „cheesey“ limestones (MONNINGER 1979). The unit includes a conspicuous amount of interbeds that consist of various, mostly light-colored shales. The dolostones are locally rich in tepee-structures, whereas algal structures remain sparse. Carbon isotope values for the Lie de vin Formation at Tiout range from -3.3 to -1.5‰ according to RENNER (1994). Similarly, MALOOF et al. (2005) indicate that the d13C values range between -3.5 and 0.5‰ for most of the Lie de vin Formation. However, the latter authors detected a peak value of 3.5‰ in the uppermost Lie de vin and lower Igoudine Formations. Nevertheless, RENNER (1994) suggested that the values identified at Tiout differ strongly from the signatures of the formation in more easterly section, such as in the Agdz area. Other
authors have correlated the Tiout section with the Zhurinsky Mys section on the Lena River of the Siberian Platform based on those data and on a few measurements of the magnetic polarities (KIRSCHVINK et al. 1991). Igoudine Formation Although the Igoudine Formation (Fig. 3: D–E) is entirely dominated by massive limestones, it can be subdivided into two members, which clearly differ in facies. The lower, informal member (about 185 m in the section) consists of limestones that indicate restricted marine conditions, but the upper Tiout Member (about 115 m) is characterized by oolitic, intraclast-bearing, or fossil hash limestones that indicate stronger wave-action. The formation commences with light and dark, thickbedded micritic limestones and dolomitic limestones with large chert nodules. The limestones are frequently laminated and grade into conspicuously laminated dolostones that locally have tepee-structures and chert nodules. Shale interbeds are rare. The total thickness of this „limestone and dolostone facies“ is about 80 m. The overlying subunit (termed the „black limestone facies“) consists of 70 m of massive, locally fetid, homogeneous black limestones with a few dark chert nodules. Shaly interbeds are almost absent. Columnar stromatolites (Acaciella fsp. cf. A. angepena, Tioutella bouddai) were found in three successive horizons near the middle of the subunit (SCHMITT 1979a, 1979b; Fig. 3: beds 63-65). The upper part shows small algal build-ups (Fig. 3: beds 66 and 67) and a few simple trace fossils. This subunit includes a notable positive d13C excursion (excursion B of LATHAM & RIDING 1990), which has been correlated with the lower Atdabanian of the Siberian Platform (KIRSCHVINK et al. 1991). The overlying 40 m of thick-bedded, light grey laminated dolostones constitute the „teepee-dolostone fa-
56 Fig. 6. Lower Igoudine Formation, „teepee-dolostone facies,“ Tiout section. Laminated dolomitic carbonates with peloids are frequently microbialites, often developed as antiformal structures. Oblique diagenetic fabrics are „immature“ tepee-structures and synsedimentary fractures. The lower part of the photo shows vugs resulting from dissolved gypsum crystals. Pen for scale.
cies“ (MONNINGER 1979) or „pale dolostone facies“ (SCHMITT 1979) (Fig. 6). The name refers to the frequent occurrence of tepee-structures, which show variable shape and size. Shale intercalations are sparse, but the facies is characterized by abundant argillaceous dolosiltites, which are almost absent in the subunit below. Interlaminations with limestones can be seen locally. The „black limestone facies“ (MONNINGER 1979) or the „dark limestone facies“ (SCHMITT 1979a) has a thickness of about 30 m. Black limestones with bed thickness of up to a few meters dominate the unit. Dolomitic limestones with thin shale intercalations occur at its base, whereas shales disappear completely higher in the section. The highest thrombolites build-ups are found towards the top (Fig. 3: beds 66-68). Recent investigations have produced Hyolithellus tubes and poorly preserved, probable trilobite shell fragments from this subunit. These undeterminable fragments are the oldest trilobite remains known from Morocco and probably from entire West Gondwana. T i o u t M e m b e r : The lower part of the Tiout Member („transition layers“ of MONNINGER 1979, and SCHMITT 1979a) consists of 20 m of „black oolitic limestone facies“ (MONNINGER 1979). The generally black, finely crystalline limestones, which are characteristic of the subunit, are either oolitic or rich in intraclasts or oncoids. The interbeds are variably developed as olive, arenitic siltstones to mottled limestones. The base of the member is relatively sharp and defined by the lowest layer of arenitic calcareous siltstones. The oldest shelly fossils described so far from Morocco are trilobites from mottled limestone interbeds close to the base of the subunit (SDZUY 1978). About 13 m higher, the first archaeocyathans occur in dark limestones (DEBRENNE & DEBRENNE 1976, 1978, 1995; SDZUY 1978, 1981).
Archaeocyathans become more frequent higher in the section in a rather distinct facies of the upper part (about 90 m) of the Tiout Member. This subunit, characterized by dark limestones with local intercalations of argillaceous and calcareous shales, is termed the „archaeocyathid-bioherm facies“ (MONNINGER 1979). Dark to lightcolored biohermal archaeocyaths-bearing limestones may laterally grade into dark, oolitic limestones. The shales are primarily developed as relief-fillings; distinct bedinterbed-alternations are not developed. The trilobites occur almost immediately above the conspicuous facies change that marks the base of the Tiout Member and have been noted in numerous higher layers (Figs. 3, 8). Thus, their occurrenc has been linked to a change to favorable habitats and does not reflect the evolutionary origin of trilobites in the area. However, shell fragments and the Hyolithellus tubes mentioned above have been found as insoluble residues of limestones in the lower member of the Igoudine Formation. These occurrences indicate that the sudden appearance of shelly fossils in the Tiout Member is most probably a preservational artifact. Most of the trilobite taxa found in the Tiout Member are primitive bigotinoids (SDZUY 1978, 1981; G. GEYER, unpubl. data, Fig. 8). Small fallotaspidids (several species of the genus Eofallotaspis) occur less frequently. The faunas have been studied by SDZUY, but few data were published. A preliminary report (SDZUY 1978) established the new bigotinid genus Hupetina with the type species H. antiqua from his lowest trilobite-bearing horizon T 1 and the new fallotaspidid genus Eofallotaspis with the species E. tioutensis (from T 10) and E. prima (from T 4). However, Eofallotaspis, among others, is known from a number of horizons (e.g., T 4, T 9, T 10, T 12, T 14, and T 17) (Fig. 8). Lithostratigraphic and biostratigraphic correlation of these layers with equivalent strata of the
57
Fig. 7. Upper Igoudine Formation, Tiout section. Mottled, aggregate-rich dark limestones as a typical facies of the Tiout Member. Scrumbly habit results from mineralized microbial mats. Photo shows original location of SDZUY’s (1978) lowermost trilobite-bearing horizon (T 1) in the section. Trilobite sclerites are best isolatable from marly, limonitic portions at the contacts of the massive limestone beds by special preparation techniques. Circular hole on the right is from KIRSCHVINK’s magnetostratigraphic investigation.
Tazemmourt section proves the younger age of most of these layers, compared with the Fallotaspis tazemmourtensis Zone (SDZUY, unpubl.), and Hupetina antiqua has been found in the Tazemmourt section in an equivalent position (SDZUY 1978). Eofallotaspis has, thus, been designated as an index genus of the Eofallotaspis Zone (GEYER 1990a). The oldest Fallotaspis species of the Tiout section (the index species Fallotaspis tazemmourtensis) has been found at T 14. The archaeocyathans have been studied in detail by DEBRENNE & DEBRENNE (1978). Additional remarks are found in DEBRENNE et al. (1992) and DEBRENNE & DEBRENNE (1995). Amouslek Formation 260 m of the lower Amouslek Formation are exposed at Tiout. Black, locally oolitic limestones alternate with thick units of greenish or grey, rarely purple, shales. Archaeocyathans are frequently found in the limestone beds, especially in the lower part, and form local bioherms. The shales consist of argillaceous and calcareous siltstones with variable carbonate content. The base of the formation is a probable unconformity marked by a thin, lenticular volcanic ash with reworked clasts of Igoudine limestone. The lower part of the formation is ca. 10 m thick complex of hard, grey shales, overlain by an alternation of limestones and shales. In contrast, the upper part of the section is dominated by shales, and the few limestones form distinct marker beds. The biota of the Amouslek Formation at Tiout does not allow for a finely differentiated biostratigraphiy. Only the Daguinaspis Zone can be recognized with certainty at the top of the exposed sequence. However, lithologic correlation with the nearby section at Tazemmourt (see below) permits a bracketing of the underlying Choubertella and Fallotaspis tazemmourtensis Zones. That part
Fig. 8. Generalized lithologic sequence of the Tiout Member with trilobite-bearing (prefix T) and archaeocyath-bearing horizons (prefix A) indicated. Sketches indicate (mostly undescribed) trilobites from the horizons studied by SDZUY. Modified from SDZUY (1978: Fig. 2b).
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of the section representing the Choubertella Zone is presumably lacking because of two faults that cross the section. LIÑAN & SDZUY (1978) noted the presence of Lemdadella tioutensis and trilobites tentatively assigned to Bigotinops from the lowest „thick limestone above the base“ of the formation (see Fig. 8). Although no index fossil is known from these beds, the authors assigned the fauna to the Fallotaspis tazemmourtensis Zone and suggested the potential of the genus Lemdadella as an index fossil to the zone. However, it should be noted that Fallotaspis tazemmourtensis occurs in the shale unit at T 14 of the Tiout Member. Cyclicity, facies, and depositional environments The cyclicity, which is visible in the Lower Cambrian sequence of the Souss Basin, is especially well illustrated by the Lie de vin through Amouslek Formations of the Tiout section (MONNINGER 1979; SCHMITT, 1979a; TUCKER 1986b; GEYER 1989b). Cyclothems and smaller-scale cycles and rhythms are recognizable in the Lie de vin and Igoudine Formations as second to fourth order cycles (Fig. 9). Simple parasequences (rhythms or cycles) of the Lie de vin Formation are a few meters in thickness and consist of simple lithosequences. They may have been primarily controlled by changes in terrigenous input. Those cycles compose units tens of meters in thickness and include several lithosequences (MONNINGER 1979). They may record variations in terrigeneous input and environment or climate and hence record changes in water depth and salinity. These cycles in turn compose ‘supercycles’ that are hundreds of meters in thickness. The latter are interpreted to be a response to regional trends in epeirogenic basin development or eustatic fluctuations. The supercycles show a long-term trend which reflects either epeirogeny or trends in eustatic changes. A changing input of terrigenous detritus and environmental changes may also be responsible for the changing lithologies represented in the section. Each cycle starts and ends with a comparatively high content of terrigeneous material, whereas the rock types in the middle of the cycles are more-or-less pure carbonate and document limited detrital input (Fig. 10). On a smaller scale, the same is recognizable for the change from dolostones though dolomitic limestones to pure limestones, a trend which also corresponds to an increase in bed thickness and is paralleled by more diverse algal build-ups (SCHMITT 1979a). Remarkably, the majority of cycles in the Lie de vin Formation are symmetrical, which suggests very low wave energies and a fluctuating balance between sedimentation rate and basinal subsidence (MONNINGER
1979). Clearly asymmetrical, tide-dominated cycles with a vertical succession that includes shaly calcareous marlstone — limestone — dolomitic marlstone — shaly calcareous marlstone are developed in the upper sequence of the Lie de vin Formation. These modal cycles show variations that can be explained by lateral transgressive–regressive environmental trends (MONNINGER 1979), but may also mark geographically widespread, coeval, aggradational/shoaling parasequences. MONNINGER (1979) and SCHMITT (1979a) related the supercycles to the relative volume of calcimicrobial buildups parts of the Lie de vin and Igoudine Formations (Fig. 9). In most cases, the relative volume of build-ups reaches its maximum when the cycle thickness starts to decrease (MONNINGER 1979). Stromatolites are limited to regressive half-cycles, but absent in transgressive halfcycles (where tepee-structures occur frequently). Cyanobacterial build-ups are additional indicators of secular facies changes. Thus, a succession consisting of a sequence including low stromatolite biostromes — thrombolites — branched columnar stromatolites — low bulbous build-ups tracks increasingly restricted facies (MONNINGER 1979; SCHMITT 1979a). In summary, the Lie de vin Formation at the Tiout section displays a remarkable, extremely shallow, low energy, schizohaline, microtidal environment with surprisingly regular changes in depositional processes. This unusual facies culminates in tidal flat dolostones with cyanobacterial laminites and evaporites now recorded by pseudomorphs after halite and gypsum, which indicate hypersaline depositional conditions. The facies contrasts generally increases up-section and are best seen by the development of more and more lithologies. Purple shales, which are intercalated in the upper Lie de vin Formation where a regressive trend is clearly indicated, have been interpreted to suggest changes in water chemistry. Geochemical data on this interval were interpreted to suggest fresh water influx (MONNINGER 1979), although the change in shale color may simply indicate a reduction in sediment accumulation rate and production of reddish shales with long residence time at the sediment–water interface. An increasing amount of carbonate from the Lie de vin Formation to the upper part of the Igoudine Formation and a synchronous decrease of detritic material are explained by growing distances from the shore with widening of the basin due to transgression which locally and temporarily led to less restricted marine conditions. At the base of the Tiout Member, the change in microfacies from argillaceous dolosiltites and dark, micritic, often finely laminated limestones and dolostones to massive oolites, pack/wackestones, and marlstones with abundant intraclasts indicates higher wave or tidal
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Fig. 9. Cyclicity in the Tiout section. The figure illustrates variations in lithology, bed thickness, bioherm height, and relative bio-herm volume from the upper Adoudou through lower Amouslek Formations. Modified from MONNINGER (1979) and GEYER (1989b).
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Fig. 10. Lithological elements that indicate cyclicity. Gradual magnification exhumes different orders of cycles. Lithologies: L, dark limestone; LD, dolomitic limestone; D, dolostone; A, argillaceous dolosiltite, DA, dolomitic marlstone; S, shale. Modified from MONNINGER (1979: Fig. 11).
energy, most probably connected with increasing water depths and more normal marine conditions. It remains unclear whether this change is created by an intracontinental event, regional climatic change, or epeirogenic or eustatic changes. Fossil record The following list indicates the fossil content of selected horizons. The prefix TIOU-M refers to the num-bering of sample horizons by MONNINGER (1979), TIOU-DB to the archaeocyathan horizons of DEBRENNE & DEBRENNE (1978), TIOU-T to the trilobite sample horizons of SDZUY (1978), and TIOU-EB to sample horizons of E. BERNEKER (unpublished). A revised taxonomy of the archaeocyathans is in DEBRENNE & DEBRENNE (1995), where a complete list of the known archaeocyathans is given. Amouslek Formation Tt-M959: Daguinaspis Zone. Sponge spicules, trilobites, hyoliths, Renalcis sp., Girvanella sp., Epiphyton sp., Kordephyton sp., undeterminable calcareous algae. Tt-M958, TIOU-1, Tt-3: Daguinaspis Zone. Microschedia amphitrite GEYER 1994, Brevipelta chouberti GEYER 1994, Daguinaspis abadiei HUPÉ 1953, Daguinaspis sp., Marsaisia robauxi HUPÉ 1953, Resserops sp., hyoliths, numerous small Planolites tubes.
TIOU-EB I/44a: Archaeocyathans, Renalcis sp., Epiphyton sp. Tt-M896: Trilobites. Tt-M883: Trilobite hash. TIOU-SZI (~Tt-M869): Brevipelta sp. cf. B. chouberti G EYER 1994, Fallotaspis sp., hyoliths, trace fossils. TIOU-M869: Chancelloriids. TIOU-SZIa (~TIOU-M867): Archaeocyathans, lingulids, trilobites, hyoliths, Hyolithellus sp. cf. H. micans BILLINGS 1872. TIOU-~M860: Probably Fallotaspis tazemmourtensis Zone. Lemdadella tioutensis L IÑAN & S DZUY 1978, Bigotinops? sp. TIOU-M856: Hyolithellus? sp. TIOU-T 17: Eofallotaspis n. sp. Tiout Member TIOU-DB 24d (TIOU-M849): Obolellid brachiopods. TIOU-DB 23: Erismacoscinus primus (D EBRENNE & D EBRENNE 1978), Protopharetra taissensis (D EBRENNE 1958), Nochoroicyathus crassus (DEBRENNE 1961), Agastrocyathus gregarius DEBRENNE 1961, Neoloculicyathus magnus DEBRENNE, Protopharetra circula (D EBRENNE 1964), Erismacoscinus fasciola (DEBRENNE & DEBRENNE 1978), Retecoscinus minutus (D EBRENNE 1959), Tumulifungia marocana DEBRENNE 1978. TIOU-DB 21: Protopharetra taissensis (DEBRENNE 1958), Protophareta sp. aff. P. grandicaveata VOLOGDIN, Protopharetra circula (DEBRENNE 1964). TIOU-DB 20: Afiacyathus compositus DEBRENNE 1961, Protopharetra circula (DEBRENNE 1964). TIOU-T 14, EB I/42a: Eofallotaspis sp., Fallotaspis tazemmourtensis HUPÉ 1953, Renalcis sp., Epiphyton sp. TIOU-T 13x (TIOU-M842): Brachiopods, trilobites.
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TIOU-DB 19 (TIOU-M842): Erismacoscinus primus (DEBRENNE & DEBRENNE 1978), Erismacoscinus fasciola (DEBRENNE & D EBRENNE 1978), Protopharetra spp., Neoloculicyathus magnus DEBRENNE & DEBRENNE, Tumulifungia marocana DEBRENNE & DEBRENNE 1978, Hyolithellus sp. cf. H. micans BILLINGS 1872. TIOU-T 12, DB 18, EB I/37: Neoloculicyathus magnus DEBRENNE & DEBRENNE, Tumulifungia marocana DEBRENNE 1978, Eofallotaspis sp., Renalcis sp. TIOU-DB 17: Neoloculicyathus magnus DEBRENNE & DEBRENNE 1978, Tumulifungia marocana DEBRENNE & DEBRENNE 1978, Afiacyathus compositus (DEBRENNE 1961), Nochoroicyathus sp., Protopharetra circula (D EBRENNE 1964), Erismacoscinus fasciola (DEBRENNE 1978), Retecoscinus minutus (DEBRENNE 1959), Dictyocyathus stipatus (D EBRENNE 1964). Tt-B17 (base of TIOU-M842?): Archaeocyathans, Hupetina sp. cf. H. antiqua SDZUY 1978, Hyolithellus sp. cf. H. micans BILLINGS 1872. TIOU-EB I/35: Archaeocyathans, Renalcis sp. TIOU-EB I/32: Archaeocyathans, sponge spicules, trilobites, Renalcis polymorphum (MASLOV), „Girvanella“ sp. TIOU-EB I/31: Archaeocyathans, sponge spicules, chancelloriids, trilobites, Epiphyton sp. TIOU-DB 15-16, EB I/30: Erismacoscinus primus (DEBRENNE & D EBRENNE 1978), Neoloculicyathus magnus DEBRENNE & DEBRENNE, Rotundocyathus sp., Protopharetra circula (DEBRENNE 1964), Erismacoscinus fasciola (D EBRENNE 1978), Tumulifungia marocana D EBRENNE & D EBRENNE 1978, Dictyocyathus stipatus (DEBRENNE 1964), Afiacyathus pruvosti DEBRENNE 1964, Agastrocyathus gregarius DEBRENNE 1961, trilobites, Renalcis sp. TIOU-T 10: Eofallotaspis tioutensis SDZUY 1978 (type locality and stratum). TIOU-DB 14, EB I/28: Erismacoscinus primus (D EBRENNE & DEBRENNE 1978), Afiacyathus compositus (DEBRENNE 1961), Neoloculicyathus magnus DEBRENNE & DEBRENNE , Protopharetra sp. aff. P. grandicaveata VOLOGDIN, Archaeopharetra sp., Tumulifungia marocana DEBRENNE & D EBRENNE 1978, Renalcis sp. TIOU-DB 13: Afiacyathus compositus (DEBRENNE 1961), Agastrocyathus gregarius DEBRENNE 1961, Tumulifungia marocana DEBRENNE & DEBRENNE 1978. TIOU-DB 12: Erismacoscinus primus (D EBRENNE & DEBRENNE 1978), Ajacicyathellus? sp., Afiacyathus compositus (DEBRENNE 1961), Agastrocyathus gregarius D EBRENNE 1961, Protopharetra circula (DEBRENNE 1964), Nochoroicyathus crassus (DEBRENNE 1961). TIOU-T 9, EB I/24: Brachiopods, Eofallotaspis sp., Renalcis sp. cf. R. novum, Kordephyton sp., Girvanella sp.
TIOU-DB 11: Erismacoscinus primus (D EBRENNE & D EBRENNE 1978), Erismacoscinus fasciola (DEBRENNE 1978), Rotundocyathus sp., Neoloculicyathus magnus DEBRENNE & D EBRENNE. TIOU-DB 10: Protopharetra taissensis (DEBRENNE 1958), Afiacyathus sp. TIOU-EB I/18: Erismacoscinus sp., spicules. TIOU-DB 1, EB I/17: Erismacoscinus primus (D EBRENNE & DEBRENNE 1978), Protopharetra taissensis (DEBRENNE 1958), Afiacyathus compositus (DEBRENNE 1961), Agastrocyathus gregarius DEBRENNE 1961, Neoloculicyathus magnus DEBRENNE & DEBRENNE, Protopharetra sp. aff. P. grandicaveata VOLOGDIN, Protopharetra circula DEBRENNE 1964, Archaeopharetra sp., Tumulocyathidae, Renalcis sp. TIOU-DB 2, EB I/16: Rotundocyathus sp., Erismacoscinus primus (DEBRENNE & D EBRENNE 1978), Afiacyathus compositus (DEBRENNE 1961), Neoloculicyathus magnus DEBRENNE & DEBRENNE, Agastrocyathus gregarius DEBRENNE 1961, Renalcis sp. TIOU-EB I/15: Renalcis sp., Girvanella sp. TIOU-DB 3: Afiacyathus sp. cf. A. compositus (DEBRENNE 1961), Pro-topharetra taissensis (DEBRENNE 1958). TIOU-DB 4: Protopharetra taissensis (DEBRENNE 1958), Girvanella sp. TIOU-DB 5: Chancelloriids. TIOU-T 5 (TIOU-M800): Brachiopods. TIOU-T 4, DB 7: Erismacoscinus primus (DEBRENNE & DEBRENNE 1978), Erismacoscinus sp., Eofallotaspis prima SDZUY 1978 (type locality and stratum). TIOU-DB 8, EB I/1 (TIOU-M797): Lowest archaeocyathans: Nochoroicyathus cribratus (DEBRENNE & DEBRENNE 1978), Erismacoscinus primus (DEBRENNE & DEBRENNE 1978), Erismacoscinus sp.; trilobite hash, „Renalcis“ sp. TIOU-T 3 (TIOU-M795): Trilobites. TIOU-T 2 (TIOU-M795): Trilobites. TIOU-T 1 (TIOU-M790): Lowest determinable trilobites: Hupetina antiqua SDZUY 1978. Lower member of Igoudine Formation TIOU-M760: Shell fragments, probably trilobite sclerites. TIOU-M755: Burrows. TIOU-M751: Shell fragments, probably trilobite sclerites. Tt-M747: Worm tubes, Hyolithellus sp. cf. H. micans B ILLINGS 1872. Tt-M724: Hyolithellus sp. cf. H. micans BILLINGS 1872. Lie de vin Formation Small burrows and burrow mottling are known from a considerable number of marly siltstones (e.g., Tt-M687).
December 2
STOP 2. TAZEMMOURT SECTION Setting
Introduction and summary
The Tazemmourt section is south of the village of Tazemmourt, about 10 km SSE of Taroudannt on the Taroudannt map sheet (Lambert coordinates 171.4/ 382.4). It lies on a marked northerly prolongation of the Anti-Atlas. The rocks dip northward, and the top of the section at Tazemmourt is covered by the alluvials of the Souss plain/valley (Fig. 1). The base of the section about 2 km south of Tazemmourt is in an east-west striking valley. The Igoudine Formation is the lowest exposed unit. The conspicuous transverse ridge north of that valley is formed by limestones that lie about at the level of ABADIE’s (1949) unit 17.
Due to its accessibility, the Tazemmourt section received a great amount of study in the first period of research on the Cambrian of the Anti-Atlas. It was the first section studied by J. ABADIE in 1949. He measured the section, which was first published by HUPÉ (1953: Fig. 7) together with a lithologic log. HUPÉ (1953, 1959) completed more detailed studies on the trilobite faunas, which had first been dealt with in a preliminary paper by HUPÉ & ABADIE (1950) which marks a starting point for a somewhat detailed Lower Cambrian biostratigraphy not only for Morocco but on a global scale.
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In HUPÉ’s first monograph (1953), a surprisingly high number of holotypes were chosen from the Tazemmourt section, and include those for such important species as Fallotaspis tazemmourtensis, F. typica, F. longa, F. longispina, F. planospinosa, Choubertella lata, C. spinosa, Daguinaspis ambroggii, D. latifrons, Pararedlichia pulchella, P. rochi, P. subtransversa, Resserops brevilimbatus, Marsaisia parvifrons, and Bigotinops dangeardi. The archaeocyathans have been described in detail by DEBRENNE (1959, 1960a, 1964; ROZANOV & DEBRENNE 1974; DEBRENNE in DESTOMBES et al. 1985), and a number of forms were first described from the section. Nevertheless, the stratigraphic interval that is exposed in the section is fairly limited and includes the Tiout Member of the Igoudine Formation and most of the Amouslek Formation. Difficulties arise from uncertainties of the published lithologic logs. Published sections (ABADIE in HUPÉ 1953; HUPÉ 1959; HOLLARD 1985) differ from each other in a number of details, especially in the lower third of the section. The section shown on Fig. 11 is mainly based on the section figured by HUPÉ (1959) but includes additional, unpublished information. Lithologic sequence The exposed sequence includes most of the Igoudine and Amouslek Formations. The sequence consists mainly of an alternation of limestones, dolostones, and shale units. The basal part includes the upper Tiout Member and is the best example of that member aside from the Tiout section. It contains a number of archaeocyathan-bearing beds, which were studied in detail by DEBRENNE (see references above and DEBRENNE & DEBRENNE 1995). In addition, trilobites (assigned to the Eofallotaspis Zone) occur as well. The shales are generally green in the lower part of the section, with purple portions in units 20 through 28. Further up-section, the shales are less distinctly green and become progressively paler. Relatively pure limestones are found in the basal part of the section. Rubbly, „scoriaceous“ limestones are intercalated as a series of intervals from units 17–29 and units 48–56. These nodular intervals represent diagenetic growth of carbonate along shell hash and siliciclastic layers, as well as burrowed layers. Biohermal limestones with archaeocyathan-algal associations are most frequently found at the top of the section (units 60, 62, and 64) and record a much different biofacies than those from the lower bedded carbonates. Dolomitic limestones and dolostones are especially frequent around 100 m below the top of the section. The list of trilobites follows HUPÉ (1953, 1959; nomenclature partly revised by GEYER et al. 1995), and is supplemented by data from SDZUY (1978) and unpub-
Fig. 11. Generalized lithologic sequence of the Tazemmourt section, based on data from ABADIE in HUPÉ (1953) and HUPÉ (1959). Fossil „sample“ horizons (prefix F) refer to HUPÉ’s (1953, 1959) levels. Numbers without prefix indicate the units of ABADIE in HUPÉ (1953) mentioned in the text. See Fig. 18 for legend.
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lished data. The nomenclature of the archaeocyathans is from DEBRENNE & DEBRENNE (1995). 283 m to top: Green shales. F 11: Antatlasia hollardi HUPÉ 1953, Antatlasia bourgini HUPÉ 1953 (type locality and probably type stratum). F 10: Fallotaspis typica HUPÉ 1953 (type locality and probably type stratum), Antatlasia hollardi HUPÉ 1953, Resserops (Resserops) brevilimbatus HUPÉ 1953, Marsaisia sp. cf. M. robauxi HUPÉ 1953. 275-283 m: Limestone beds with archaeocyathans. 242-275 m: Three limestone complexes with archaeocyathans, separated by greenish slaty shales. Upper shale unit (63) yields obolellid brachiopods, the limestone units (60 and 62) have the archaeocyathans (DB 20) Mennericyathus echinus (DEBRENNE 1964), Protophareta gemmata (DEBRENNE 1964), Afiacyathus pruvosti DEBRENNE 1964, and Dictyocyathus stipatus DEBRENNE 1964. 228-242 m: Green, slaty shales. Upper part of shales (unit 59) with F 9: Marsaisia robauxi HUPÉ 1953, Resserops uncioculatus HUPÉ 1953. Lower part of shales (unit 57) with F 8: Resserops uncioculatus HUPÉ 1953. 225-228 m: Thick beds of rubbly limestone (Calcaire scoriacé). 218-225 m: Green shales with bed of Calcaire scoriacé at the base. Shales (unit 55) with F 7: Daguinaspis ambroggii HUPÉ & ABADIE 1950, Marsaisia parvifrons HUPÉ 1953. 214-218 m: Green shales, with Calcaire scoriacé bed at the base. Shales (unit 53) with F 6: Daguinaspis ambroggii HUPÉ & ABADIE 1950 (type locality and probably type stratum), Daguinaspis latifrons HUPÉ 1953 (type locality and type stratum), Resserops falloti HUPÉ 1953. 211-214 m: Green shales, with Calcaire scoriacé bed at the base. Shales (unit 51) with F 5: Brevipelta chouberti GEYER 1994, Daguinaspis ambroggii HUPÉ & ABADIE 1950, D.
Fig. 12. Daguinaspis ambroggii Hupé 1953, cranidium, holotype, Tazemmourt section, most probably from unit 53. Musée Nationale d’Histoire Naturelle, Paris.
subabadiei HUPÉ 1953 (type locality and type stratum), Marsaisia robauxi HUPÉ 1953. TAZE-CH 11: Brevipelta chouberti GEYER 1994, Daguinaspis abadiei HUPÉ 1953, D. ambroggii HUPÉ & ABADIE 1950, Marsaisia robauxi HUPÉ 1953, Resserops (Resserops) bourgini HUPÉ 1953, burrows. 204-211 m: Green shales, with Calcaire scoriacé bed at the base. Shales (unit 49) with F 4: Daguinaspis abadiei HUPÉ 1953 (type locality and type stratum), Marsaisia parvifrons HUPÉ 1953 (type locality and type stratum), Resserops brevilimbatus HUPÉ 1953 (type locality and type stratum). Archaeocyathans (DB 18): Neoloculicyathus abadiei (D EBRENNE 1959), Mennericyathus echinus (DEBRENNE 1964), Erismacoscinus marocanus DEBRENNE 1958. 185-204 m: Alternation of archaeocyathan-bearing limestones, Calcaire scoriacé beds, and green shales. 174-185 m: Complex of massive, dolomitic limestones. 159-174 m: Green shales, grading up-section into alternation of limestones and shales. 139-159 m: Alternation of archaeocyathan-bearing limestones, Calcaire scoriacé beds, and shale layers. Shales of lower 15 m unfossiliferous. Slaty shales of unit 33 in the upper part represent F 3: Archaeocyathans, Fallotaspis longispina HUPÉ 1953 (type locality and probably type stratum). 134-139 m: Green and purple shales, at the base with a lenticular massive Calcaire scoriacé bed (unit 31) termed „Barre de Ksars“ (compare DEBRENNE & DEBRENNE 1995). Shales of unit 28 represent F 2: Brevipelta chouberti GEYER 1994, Fallotaspis longa HUPÉ 1953, Fallotaspis planospinosa HUPÉ 1953 (type locality and probably type stratum), Choubertella crassioculata HUPÉ 1953 (type locality and type stratum), Choubertella lata HUPÉ 1953 (type locality and type stratum), Choubertella spinosa HUPÉ 1953 (type locality and type stratum), hyoliths. 117-134 m: Green shales, with purple shales at the top, alternating with Calcaire scoriacé beds. Trilobites unknown. Archaeocyathans (DB 13) include: Retecoscinus minutus (DEBRENNE 1959), Protopharetra gemmata (DEBRENNE 1964), Neoloculicyathus abadiei (DEBRENNE 1959), Tumulocoscinus equiporus (DEBRENNE 1959), Erismacoscinus marocanus (DEBRENNE 1958). 61-117 m: Alternation of thick units of greenish shales, with purple shales at the top, archaecyathan-bearing limestone beds of 1 to 1.5 m thickness at the base of each unit. Trilobites are found throughout the shales and are dealt with as bulk sample F 1 by HUPÉ (1959): Fallotaspis tazemmourtensis HUPÉ 1953 (type locality and type stratum), Fallotaspis longa HUPÉ 1953 (type locality and probably type stratum), Fallotaspis plana HUPÉ 1953, Fallotaspis planospinosa HUPÉ 1953, Pararedlichia pulchella HUPÉ 1953 (type locality and type stratum), Pararedlichia rochi HUPÉ 1953 (type locality and type stratum), Pararedlichia subtransversa HUPÉ 1953 (type locality and type stratum), Bigotinops dangeardi HUPÉ 1953 (type locality and type stratum), Pruvostinoides? sp.; archaeocyathans of DB 10 and DB 12: Mennericyathus echinus (DEBRENNE 1964), Erismacoscinus marocanus DEBRENNE 1958, Tumulocoscinus equi-
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porus (DEBRENNE 1959), Nochoroicyathus crassus (DEBRENNE 1961), Protopharetra stipata (DEBRENNE 1964), Neoloculicyathus abadiei (DEBRENNE 1959), Agastrocyathus gregarius DEBRENNE 1961, Retecoscinus minutus (DEBRENNE 1959). Frequently with Brevipelta chouberti GEYER 1994. 24-61 m: Bluish-green shales with intercalated thin to very thick limestone beds. Contains rare Brevipelta sp. cf. B. chouberti GEYER 1994. 0-24 m: Tiout Member. Dark to medium-grey massive limestone beds, locally oolitic and oncolitic. TAZE-T 0: Eofallotaspis sp., Fallotaspis? sp., Bigotinops? sp. DB 4-5: Protopharetra circula (DEBRENNE 1964), Protopharetra gemmata (DEBRENNE 1964), Erismacoscinus marocanus DEBRENNE 1958, Erismacoscinus primus DEBRENNE & DEBRENNE, Rotundocyathus sp., Agastrocyathus gregarius DEBRENNE 1961, Retecoscinus minutus (DEBRENNE 1959), Nochoroicyathus crassus (DEBRENNE 1961), Neoloculicyathus abadiei (DEBRENNE 1959), Afiacyathus compositus (DEBRENNE 1961), Mennericyathus echinus (DEBRENNE 1964).
Fossils and biostratigraphy The large number of trilobite taxa based on type material from the section indicates that the concept of HUPÉ’s Zones I to III (HUPÉ & ABADIE 1950; HUPÉ 1953) was based on the faunal succession of the Tazemmourt section. HUPÉ (1959) restudied the trilobite faunas and emphasized a number of levels that helped refine the faunal sequence (HUPÉ 1953). These levels have been listed above according to their stratigraphic position. The taxa cited by HUPÉ (1959) have been revised in part by GEYER (1990b, 1995) and in GEYER & LANDING (1995), and the new taxa listed as nomina nuda by HUPÉ (1959) are omitted. However, it should be noted that the majority of HUPÉ’s types were from collections made by ABADIE, who did not record the exact sample horizon. Furthermore, a number of taxa are based on single specimens that are unfavorably preserved, and a number of taxa require revision. As an example, the three species of Pararedlichia are too poorly preserved and sparsely represented to permit recognition of their distinct characters. Hence, they cannot be used in intercontinental correlations, and their assignment to Eoredlichia (a genus which is primarily characterized by features of the thorax and pygidium that are unknown from any species of Pararedlichia) as known from the Yangtze Platform, as suggested by PILLOLA (1991), is questionable. Subsequent investigation of the earliest, Eofallotaspis to Fallotaspis tazemmourtensis Zone trilobites of the section suggests that their succession is well documented, but needs to be completely reported (SDZUY 1978).
Fig. 13. Fallotaspis tazemmourtensis Hupé 1953. Preserved is a cepalon with attached partial thorax and disarticulated rostal plate, obviously showing a carcasse emedded just after molting of the animal. Holotype, Tazemmourt section, from HUPÉ’s bulk sample F1. Musée Nationale d’Histoire Naturelle, Paris.
These oldest trilobites belong to assemblages that are approximately coeval with those from the Tiout Member of the Igoudine Formation at Tiout (TAZET0; SDZUY 1978, unpublished). Nevertheless, the section features a transition from the Eofallotaspis to the Fallotaspis tazemmourtensis Zone. Fallotaspis tazemmourtensis (Fig. 13) is reported only from HUPÉ’s bulk sample F 1 and probably occurs as high as the shale unit 20, which appears to mark the top of the zone. However, the base of the zone is not yet identified, and it is thus not clear whether the base of the Amouslek Formation belongs to the Eofallotaspis Zone or the Fallotaspis tazemmourtensis Zone. The Choubertella Zone ranges up to about unit 28, but Choubertella is represently by relatively few specimens that do not permit precise identification of the limits of the biozone. The Daguinaspis Zone includes the overlying sequence, probably up to shale unit 59 (= F 9). Unit 61 yields Antatlasia hollardi, the index fossil of the overlying A. hollardi Zone.
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December 3
LEMDAD SYNCLINE Setting The Lemdad syncline is located on the southern rim of the west-central High Atlas in the vicinity of the ‘Bloc occidental’ of the ‘Massif ancien’ (Adrar n’Dren massive or Promontoire d’Ouzellarh sensu CHOUBERT). The name „Lemdad“ is a westernization of „Oued el Mdad,“ a fairly permanent creek that crosses the area and provides the best outcrops (Fig. 14). Structurally, the region features a simple northweststriking syncline with a subvertical fault along the axis. The axis itself dips northwestward so that the outcrop pattern is v-shaped. Complete successions of lowest Cambrian through upper Middle Cambrian rocks form the core of the syncline. These Cambrian rocks are locally covered with almost flat-lying late Cretaceous sedimentary rocks which were deposited during the Cenomanian–Turonian transgression. By contrast, the western and southern parts of the flanks of the syncline show a fairly complete lowest Cretaceous through Turonian succession. The Lemdad area was earlier and is still, in part, termed ‘Ouneïn’ by various authors. However, the name Ouneïn is correctly applied to the area northeast of the Lemdad
syncline, where additional outcrops can be found, but where the sections are much more complicated by faults. The outcrop in the Lemdad syncline permits a detailed stratigraphical comparison between numerous short sections across the Lower to Middle Cambrian transition (Fig. 23) and clearly shows the conspicuous facies changes from the southwestern to the northeastern flank of the syncline. The area is thus probably the best to study facies changes across the Tatelt and lower Jbel Wawrmast Formations and further permits recognition of lateral microfacies changes of the limestone beds of the Brèche à Micmacca Member. In addition, the area has the only highly fossiliferous middle Lower Cambrian strata west of the ‘Massif ancien’. Even more importantly, the Cambrian successions in the Lemdad Syncline have been central in defining the Lower–Middle Cambrian boundary interval sequence stratigraphy of southern Morocco and in distinguishing epeirogenic vs. eustaric controls on the development of depositional sequences (LANDING et al. 2006). Briefly summarized, strong early Middle Cambrian epeirogenic uplift on fault-bounded blocks led to regional denudation
Fig. 14. Sketch map of the Cambrian of the Lemdad syncline with outcrops through the upper Tata Group (brick pattern), the Feijas internes Group (densely stippled), and location of sections Le I to Le XVI. Pattern at Jbel Tousham indicates outcrops of Cretaceous rocks (see Stops 3 and 7). Rivers are in black, roads in grey.
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and erosive loss of the Tazlaft Formation in the High Atlas and to erosion as low as the Pan-African orogen further to the north. Stops at sections Le I–IIA, Le IIB, and Le XI show evidence for an unconformity on the Issafen Formation marked by a thin polymict conglomerate with clasts of Pan-African vein quartz and welded red rhyolite at the base of the Tatelt Formation and a progressive, west–east, ca. 52 m erosive cutout of the upper Issafen Formation. The Tatelt Formation marks a regional marine inundation of the Moroccan Atlas regions that was driven by epeirogenic processes demonstrated by the high volcanogenic component of the Tatelt Formation. Regional offlap after deposition of the Tatelt led to its regional erosion and development of a sequence boundary on the Tatelt. The depth of this erosion was least in the Lemdad Syncline, as the Lemdad area is the
only region where the Tatelt Formation extends as high as the Cephalopyge Zone (section Le II and Le IX). In the Anti-Atlas, the Tatelt Formation is eroded as low as the Hupeolenus Zone. Epeirogenic basin reorganization within the Cephalopyge notabilis Chron led to foundering of the Souss Basin and regional deposition of the Jbel Wawrmast Formation (LANDING et al. 2006). The initial deposits of the Jbel Wawrmast include condensed, fossiliferous carbonates and siliciclastic mudstones of the Brèche à Micmacca Member. The cool-water, fossil hash limestones of the Brèche à Micmacca Member mark the last significant carbonate deposition in Morocco, and bedded limestones comprise an exceptionally rare component of higher Cambrian strata in the Atlas and Anti-Atlas mountain regions (LANDING et al. 2006).
Fig. 15. Longitudinal, west–east section through the Lower Cretaceous through Turonian rock succession along the southern flank of the High Atlas. The section indicates regressional and transgressional developments as well as the migration of the litoral zone in the Cretaceous Atlas Gulf. Cross-section at bottom depicts true thicknesses in relation to the lateral extension. Modified from WURSTER & STETS (1982: Fig. 6).
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December 3
Stop 3. Cretaceous deposits at the southern flank of the Lemdad Syncline The drive along the southwestern flank of the Lemdad Syncline first crosses a thick unit of Cretaceous deposits. Transgressions and regressions in the so-called Atlas Gulf (a funnel-shaped bay that stretched from Agadir in the west to Safi in the east) controlled the development of variable facies during Early Cretaceous deposition. Considerable input of terrigenous clastics characterized regressive periods (late Hauterivian, early Aptian) whereas carbonates dominated during periods of accelerated sea-level rise (Valanginian, Barrêmian, late Aptian (BEHRENS & SIEHL 1982; WURSTER & STETS 1982). Westward progradation of siliciclastic detritus from the source areas in the Precambrian crystalline massifs of the central High Atlas is responsible for a decrease in grain-size toward the west. As a result, a distinct facies pattern and increase of thickness toward the west are easily recognized through the western High Atlas. The early Cretaceous through Cenomanian deposits reach more
than 1500 m in thickness at Agadir whereas the thickness is less than 40 m a short distance south of the Lemdad Syncline at Aoulouz (BEHRENS et al. 1978) (Fig. 15). The most distinctive transgression took place at the upper Cenomanien to initial Turonian, and accompanied a major flooding of the area (BEHRENS et al. 1978, WIEDMANN et al. 1978, WURSTER & STETS 1982, STAMM & THEIN 1982). As a result, Turonian rocks in the western High Atlas Mountains form a relatively uniform succession of carbonates that was originally termed the ‘barre turonienne’ (ROCH 1930, CHOUBERT 1948, BUSSON 1970). The section on the southwestern flank of the Lemdad Syncline shows a thickness of ca. 300 m for the Lower Cretaceous through Cenomanian rocks. Strong siliciclastic input led to deposition of the dominant redcolored sedimentary rocks. Approximately the upper twothirds of the succession (ca. 145 m) is Cenomanian in age (although biostratigraphic data are lacking and the base of the Cenomanian is determined only by means of facies changes). The lower part and most of the Cenomanian is dominated by dolomitic arenites with local trace fossils and mollusks (WURSTER & STETS 1982). Three intervals contain gypsum layers, and four intercalations of gray to yellowish dolostones and dolomitic arenites can be distinguished. The facies indicates deposition in lagoonal settings with occasional marine onlaps (WURSTER & STETS 1982). The underlying upper Albian is the only interval in this succession which forms a thick, carbonate-dominated unit. This characteristic interval was thus assigned a distinct name, and is termed the ‘Vraconian.’
Fig. 16. Typical Cenomanian–Turonian succession in the vicinity of the Lemdad Syncline. The reduced thickness from Tafinegoult (Tizi n’Test road) towards the Aouzer section is a result of the paleogeographic situation at the margin of the Atlas Gulf. Modified from WURSTER & STETS (1982, fig. 4).
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The Turonian limestones form a ca. 45 m-thick succession. Despite its apparently roughly homogenous lithology, distinct differences can be recognized between the lower and upper parts of the Turonian succession. Extremely high Mn/Fe ratios are measured in the lower 20 m, and are consistant with local Mn-ore deposits at the coastal fringes of the Atlas Gulf (STAMM & THEIN 1982). These conditions are interpreted to reflect a period of high sea-level with widely distributed reducing conditions which were responsible for the enrichment of such typical substances as organic carbon, strontium and phosphate. The manganese apparently accumulated in relatively deeper environments of the coastal shallow sea near the redox boundary. By contrast, lower sealevel during the late Turonian created highly saline conditions which led to a loss of normal marine facies, and a significant accumulation of dolomitic rocks in the upper part of the succession (STAMM & THEIN 1982).
Fig. 17. Schematic block diagrams of the scenery in the Cretaceous Atlas Gulf. The four diagrams illustrate extreme regressive (Hauterivian and Albian) and transgressive (Aptian and Cenomanian) conditions. View from southwest, highly exaggerated vertically. Modified from WURSTER & STETS (1982, fig. 10).
December 3
Stop 4. Section Le I Setting and introduction Section Le I is the most complete Cambrian section known from the High Atlas. This north-dipping sequence merits special attention because other sections with fossiliferous and biostratigraphically important early trilobitebearing Lower Cambrian strata are not known elsewhere in the High Atlas. The section crops out above the Oued El Mdad on the south-western flank of the syncline (Fig. 14) and has been the subject of study by a number of authors (BOUDDA 1968; BOUDDA & CHOUBERT 1972; BOUDDA et al. 1975, 1979; DEBRENNE & DEBRENNE 1975; SDZUY 1978, 1981; LIÑAN & SDZUY 1978; SCHMITT 1979b; GEYER 1983, 1988a, 1988b, 1989a, 1990a; SIEGERT 1986; DEBRENNE 1992a). The most complete lithologic log of the lower part of the section has been compiled by BOUDDA et al. (1975) and GEYER et al. (1995) and is refigured here (Fig. 19; slightly modified). The upper part has been summarized by GEYER (1990a) and GEYER et al. (1995).
Lithologic sequence Lie de vin Formation The sequence starts with shaly-calcareous deposits, which comprise the thin the Lie de vin Formation in the Lemdad Syncline — and form the only identified Lie de vin section in the High Atlas. The formation is truncated by a fault and thrust against Cretaceous deposits so that the base of the formation is unknown from the High Atlas. Lemdad Formation The Lemdad Formation is a heterolithic alternation of brownish to red and greenish marlstones and siltstones, with thick units of limestone and dolostone and frequent intercalations of variable volcanics or volcanoclastic material (Figs. 18, 19). Section Le I is the type locality of the Lemdad Formation (GEYER 1989a) and shows a rhyth-
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Fig. 18. Generalized lithological log of section Le I (left; after GEYER et al. 1995; modified from GEYER 1990a; based on data from BOUDDA 1975, BOUDDA et al. 1975, GEYER 1983, and unpublished data of SDZUY and SCHAER) and enlarged upper Lemdad Formation to Brèche à Micmacca Member in the Le I section (right; after GEYER et al. 1995; based on data from GEYER 1983, and unpublished data of K. SDZUY). TAT, Tatelt Formation. Numbers refer to sample horizons mentioned in the text and listed below.
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71
Fig. 19. Lithologic log of the lower part of section Le I. Sample numbers L .. and F .. refer to trilobite horizons of SDZUY (1978 and unpublished), S .. to stromatolitic horizons of SCHMITT (1979), A 1 and A 2 to ash sample horizons collected by E. LANDING (unpublished data). Modified from BOUDDA et al. (1975).
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Fig. 20. Lemdad Formation, lower member, with more-or-less massive limestone bed in the lower part of the photo, overlain by platy to nodular marly limestone-shale alternation. The massive limestone is largely built by columnar stromatolites (here: type locality and type stratum of Madiganites lemdadensis SCHMITT 1979. Lemdad Syncline, Section Le I, stratum S1 (see Fig. 19).
mic alternation of marlstones to siltstones and intercalated limestones and dolostones. The lower part of the formation is relatively rich in dark oolitic, oncolitic, and dark micritic limestone beds. Towards the top, the carbonate sequences are formed more and more by microcrystalline dolostones and light colored, partly sparitic limestones (BOUDDA et al. 1975; GEYER 1989a). The carbonate content and bed thicknesses decrease as well upsection, and a calcareous lower member may be distinguished from an essentially siliciclastic upper member. Various types of fine siliciclastic units are present but are mainly composed of unfossiliferous slaty siltstones. These are essentially brown, red or grey in the lower part and increasingly grey or green in the upper part. Thin, often nodular marlstone bands are intercalated with less frequent micaceous quartz to feldspathic sandstone interbeds. The cyclic succession is masked by intercalations of basic to intermediate volcanoclastics, which are generally intercalated with carbonate units but also occur as distinct beds within shale sequences. The lower boundary of the Lemdad Formation (roughly comparable to the unit termed „Calcaire Supérieur“ in the literature) has been matter of discussion in earlier articles (BOUDDA & CHOUBERT 1972; SDZUY 1978; BOUDDA et al. 1979). The problems focus on the stratigraphic assignment of reddish dolomitic beds that underlie the first limestones that yield shelly faunas. These beds were alternatively assigned either to the top of the „Série Liede-vin“ or the „Calcaire Supérieur ... in the Lie de Vin facies.“ In fact, the boundary between the Lie de vin and Lemdad Formations most probably does not coincide with the boundary between the Lie de vin Formation and the Igoudine Formation elsewhere, and the layers in question are secondarily red-colored and belong to the Lemdad Formation. Tied to this problem is the claim of
trilobites found at the top of the Lie de vin Formation (BOUDDA & CHOUBERT 1972). Trilobite-bearing horizons are found throughout the section and start close to the base of the Lemdad Formation (Figs. 18, 19). As mentioned above, the first trilobites in the section are found more-or-less immediately above the rocks of Lie de vin-type facies (i.e., the reddish colored strata of the basal Lemdad Formation). The trilobites were partly studied by SDZUY, and preliminary results are presented in SDZUY (1978, 1981). However, most trilobites from the lower part of the section are still undescribed. The lower horizons are characterized by the genus Lemdadella, which was named from this locality (SDZUY 1978) and subsequently described from the Tiout section and from southern Spain (SDZUY & LIÑAN 1978). The range of the genus, however, extends relatively far upsection, and the stratigraphic range of the genus is not known with certainty. Therefore it is questionable whether Lemdadella characterizes a biostratigraphic zone with a range of the genus comparable to that established in the Anti-Atlas. Other trilobite taxa that occur in the lower Lemdad Formation include such bigotinoid trilobites (SDZUY 1981) as Bigotina and Bigotinops, the range of which needs further examination. Nevertheless, LIÑAN & SDZUY (1978) assigned the layers to the Fallotaspis tazemmourtensis Zone because of the presence of a Lemdadella species known at the base of the Amouslek Formation in the Tiout section (see above, Stop 1). The only biostratigraphically distinct datum that links the lower part of section Le I with the Anti-Atlas is the occurrence of Daguinaspis in horizon L 19, roughly 200 m above the base of the Lemdad Formation (SDZUY 1978), where a correlation with the Daguinaspis Zone is corroborated by the co-occurrence of Pruvostina sp.
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Columnar stromatolites are suprisingly frequent in the lower, trilobite-bearing part of the Lemdad Formation. Two new form-species, Madiganites lemdadensis and Vetella sarfatiae, were described from these beds (SCHMITT 1979a, 1979b; see Fig. 20). In contrast, archaecyathans occur first in the probable Daguinaspis Zone but remain frequent faunal elements up to the Sectigena Zone. They form two conspicuous bioherm levels in the upper part of the section. These levels include lower bioherms in the probable Antatlasia guttapluviae Zone in the upper part of the Lemdad Formation and upper bioherms in the Sectigena Zone at the base of the Issafen Formation (see Fig. 12 in „Latest Ediacaran and Cambrian of the Moroccan Atlas regions,“ this volume). Archaeocyathans can be found lower in this section, but the lowest appearance of these archaeocyathans clearly postdates the earliest occurrences of the group in the Anti-Atlas. DEBRENNE & DEBRENNE (1975, 1995) and DEBRENNE et al. (1992) distinguished two archaeocyathan levels at section Le I, and a supposed third level represented at section Le XI (see Stop 7, below). The most frequent SSF is Hyolithellus sp. cf. H. micans, which is especially abundant in the archaeocyathan build-ups. Hyoliths may be enriched at the tops of limestone beds. Other faunal elements are rare. Issafen Formation The Issafen Formation in the Lemdad syncline represents a relatively offshore depositional environment of this unit. A number of isolated, calcarenitic or sandy limestone beds or pure, fine-grained sandstone layers, intercalated in greenish shales are present. The fine-grained sandstones become increasingly abundant and thicker towards the top of the Issafen Formation, and are characterized by prominent HCS bedding. These HCS sand-
Fig. 21. Lemdad Formation, upper member. Archaeocyathan bioherm. Intercalated siltstones indicate variable extend of the archaeocyathan–microbial build-ups. Marginal and incorporates limestone– marlstone layers of the bioherm include numerous trilobites. Lemdad Syncline, section Le I, horizon Le I-I/1 and Le I-I/2 („Ounein A“ of DEBRENNE et al. 1992).
stones become amalgamated in the uppermost Issafen Formation, and suggest a shoaling-up facies trend (LANDING et al. 2006). The calcarenitic beds frequently contain numerous specimens of the trilobites Sectigena spp., Hebediscus lemdadensis, and Antatlasia. However, it is problematical why the identity of the Sectigena species changes from bed to bed but does obviously does not reflect a phylogenetic development. A single clearly determinable cranidium of Protolenus (Hupeolenus) hupei GEYER 1990a and poorly preserved fragments attributed to the same species, were recovered 40 cm below the top of the Issafen Formation, and just below the polymict conglomerate that marks the base of the Tatelt Formation (LANDING et al. 2006). This findings show that the top of the Issafen Formation is referable to the lowest Middle Cambrian Hupeolenus Zone and that a distinct biostratigraphic gap is not present. As noted above, the upper bioherm level with archaeocyathans and calcareous algae is found close to the base of the formation („Ounein B“ sensu DEBRENNE et al. 1992). For lithology and depositional environment of the Tatelt and Jbel Wawrmast Formation, see explanations under section Le II below. Fossil record Upper part of Jbel Wawrmast Formation Le Ia-X 243: Probably Badulesia tenera Zone. Obolellid brachiopod, „Acrothele“ sp., Skrejaspis sp. aff. S. tosali SDZUY 1967, Ornamentaspis frequens GEYER 1990, Ctenocephalus (Ctenocephalus) n. sp., „Paradoxides“ sp., echinoderm ossicles. Le Ia-X 244: Badulesia tenera Zone. „Acrothele“ sp., Kymataspis arenosa G EYER 1990, „Paradoxides“ sp., „Parasolenopleura“ sp., Conocoryphe (Parabailiella) sp. aff. C. (P.) schmidti S DZUY 1957, Badulesia tenera (H ARTT 1868), Badulesia sp. cf. B. paschi (SDZUY 1958), echinoderm ossicles. Le I-1/45: Kymataspis arenosa Zone. „Paradoxides“ sp., Parasolenopleura lemdadensis GEYER 1998.
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Le I-1/38: Parasolenopleura lemdadensis GEYER 1998. Le I-1/30: Kymataspis arenosa Zone. Kymataspis arenosa GEYER 1990, Conocoryphe (Parabailiella) sp., Parasolenopleura lemdadensis GEYER 1998, Asturiaspis sp., echinoderm ossicles. Le I-1/25: Kymataspis arenosa Zone. Parasolenopleura? sp. Le I-1/21: Kymataspis arenosa Zone. Kymataspis arenosa GEYER 1990. Le I-1/20: Kymataspis arenosa Zone. Kymataspis arenosa GEYER 1990, Paradoxides (Acadoparadoxides) sp. Le I-1/18: Kymataspis arenosa Zone. Kymataspis sp. cf. K. arenosa GEYER 1990. Le I-1/~15: Kymataspis arenosa Zone. Kymataspis arenosa GEYER . Le I-1/12,5: Kymataspis arenosa Zone. „Acrothele“ sp., Kymataspis arenosa GEYER 1990, Paradoxides (Acadoparadoxides) sp., Parasolenopleura? sp., echinoderm ossicles. Le I-1/10: Kymataspis arenosa Zone. Ellipsocephalus sp. aff. E. polytomus (LINNARSSON 1877), Kymataspis arenosa GEYER 1990, „Paradoxides“ sp., Parasolenopleura? sp., echinoderm ossicles. Le I-1/9: Kymataspis arenosa Zone. „Acrothele“ sp., Ellipsocephalus sp. aff. E. polytomus (LINNARSSON 1877), Kymataspis arenosa G EYER 1990, Paradoxides (Acadoparadoxides) sp., Parasolenopleura? sp., Conocoryphe (Parabailiella) sp., Bailiella sp. aff. B. emarginata, echinoderm ossicles. Le I-1/8: Ornamentaspis frequens Zone. Ornamentaspis sp. cf. O. frequens GEYER 1990, Conocoryphe (Parabailiella) sp. Le I-1/7: Ornamentaspis frequens Zone. Trematobolus splendidus GEYER & M ERGL 1995, Ornamentaspis sp. cf. O. frequens GEYER 1990, Kymataspis arenosa GEYER 1990, Ellipsocephalus sp. aff. E. polytomus (L INNARSSON 1877), Acanthomicmacca sp. cf. A. neltneri HUPE 1953, Paradoxides (Acadoparadoxides) nobilis GEYER 1998, Parasolenopleura? sp., echinoderm ossicles, complete undeterminate echinoderms. Le I-1/5: Ornamentaspis frequens Zone. Paradoxides (Acadoparadoxides) nobilis GEYER 1998, Parasolenopleura? sp., Bailiella sp. aff. B. emarginata (LINNARSSON 1877), hyoliths. Le I-1/4,5: Ornamentaspis frequens Zone. „Acrothele“ sp., Paradoxides (Acadoparadoxides) sp., Bailiella sp. aff. B. emarginata (LINNARSSON 1877), hyoliths. Le I-1/4: Ornamentaspis frequens Zone. „Acrothele“ sp., Conocoryphe (Parabailiella) sp. Le I-1/3,5: Ornamentaspis frequens Zone. Paradoxides (Acadoparadoxides) sp., Parasolenopleura? sp., echinoderm ossicles. Le I-X 254 (Le I-1/3 bis 1/3,5): Ornamentaspis frequens Zone. „Acrothele“ sp., Condylopyge sp., Ornamentaspis frequens GEYER 1990, Kingaspidoides sp. aff. K. laetus GEYER 1990, „Paradoxides“ sp. aff. P. asturianus SDZUY 1968, Jincella? sp., Skrejaspis sp. aff. S. tosali SDZUY 1967, Conocoryphe (Parabailiella) sp., Bailiella sp., hyoliths, echinoderm ossicles, entire echinoderms. Le I-1/3: Ornamentaspis frequens Zone. Ornamentaspis frequens G EYER 1990 (incl. protaspis), Paradoxides (Acadoparadoxides) sp., Parasolenopleura? sp., Skrejaspis sp., ptychopariid new genus and species, echinoderm ossicles. Brèche à Micmacca Member Le I-1/2: Ornamentaspis frequens Zone. Ornamentaspis frequens GEYER 1990, Paradoxides (Acadoparadoxides) sp., Skrejaspis sp., Parasolenopleura? sp., Bailiella sp., echinoderm ossicles. Le I-2, X 253: Ornamentaspis frequens Zone. Ornamentaspis frequens GEYER 1990 (type locality and stratum), Kingaspidoides sp. aff. K. laetus GEYER 1990, Paradoxides (Acadoparadoxides) nobilis GEYER 1998, Parasolenopleura? sp., ptychopariid new genus and species, „Syspacephalus“ sp., Helcionella capula GEYER 1986. Le I-3: Ornamentaspis frequens Zone. Ornamentaspis frequens GEYER 1990. Le I-4: Lowermost Ornamentaspis frequens Zone. Trematobolus
Le
Le
Le
Le Le Le Le
splendidus GEYER & MERGL 1995, Cephalopyge notabilis GEYER 1988, Latikingaspis sulcatus GEYER 1990, Latoucheia (Latoucheia) pusilla GEYER 1990, Ornamentaspis frequens GEYER 1990, Parasolenopleura? sp. I-4A: Cephalopyge Zone. Trematobolus sp. cf. T. splendidus GEYER & MERGL 1995, Cephalopyge notabilis GEYER 1988, Latikingaspis alatus (HUPÉ 1953). I-5: Cephalopyge Zone. Trematobolus splendidus GEYER & MERGL 1995, Cephalopyge notabilis GEYER 1988, Latikingaspis alatus (HUPÉ 1953), Latoucheia (Latoucheia) pusilla G EYER 1990, Kingaspidoides? sp. C, Ornamentaspis sp. cf. O. frequens G EYER 1990, Gigoutella atlasensis H UPÉ 1953, Acanthomicmacca sp. cf. A. neltneri H UPE 1953, Paradoxides (Acadoparadoxides) sp. cf. A. nobilis GEYER 1998, Parasolenopleura? sp., Helcionella capula G EYER 1986. I-7: Cephalopyge Zone. Dictyonina? sp., Trematobolus splendidus G EYER & M ERGL 1995, Pseudocobboldia pulchra (HUPÉ 1953), Cephalopyge notabilis GEYER 1988, Kingaspidoides? sp. C, Latikingaspis alatus (HUPÉ 1953), Protolenus (Protolenus) sp., Protolenus (Protolenus) densigranulatus GEYER 1990, Latoucheia (Pseudolenus) ourikaensis (HUPÉ 1953), Latoucheia (Latoucheia) pusilla GEYER 1990 (type locality and stratum), Paradoxides (Acadoparadoxides) sp. cf. A. nobilis GEYER 1998, Latouchella sp. cf. L. comma GEYER 1986, Helcionella capula GEYER 1986, helcionellid genus and sp. C, Marocella mira GEYER 1986, various hyoliths, „Allatheca“ sp. (opercula), echinoderm ossicles, problematic SSFs. I-7B: Cephalopyge Zone. Latoucheia (Latoucheia) pusilla GEYER 1990. I-7/13: Cephalopyge Zone. Latoucheia (Latoucheia) pusilla GEYER 1990. I-8: Cephalopyge Zone. Cephalopyge notabilis GEYER 1988. I-9: Cephalopyge Zone. Latoucheia (Latoucheia) pusilla GEYER 1990, Helcionella capula GEYER 1986.
Tatelt Formation Le I-X 251: Probably Hupeolenus Zone. Trematobolus sp. cf. T. splendidus GEYER & MERGL 1995, Protolenus (Hupeolenus) hupei G EYER 1990, Protolenus (Hupeolenus) termierelloides GEYER 1990, Acanthomicmacca sp., Marocella mira GEYER 1986, problematics. Le I-10/03: Hupeolenus Zone. Trematobolus splendidus GEYER & M ERGL 1995, Protolenus (Hupeolenus) sp. cf. P. (H.) hupei G EYER 1990, Protolenus (Hupeolenus) termierelloides GEYER 1990, Protolenus (Hupeolenus) sp. A, Acanthomicmacca n. sp. A, „Paradoxides“ sp. Le I-10/07: Sectigena Zone. Sectigena sdzuyi GEYER 1990. Le I-10/09: Sectigena Zone. Trematobolus? sp., Berabichia sp. indet., Sectigena sdzuyi GEYER 1990. Issafen Formation Le I-11: Sectigena Zone. Hebediscus lemdadensis GEYER 1988, Protolenus (Hupeolenus) hupei G EYER 1990. Le I-F 13: Sectigena Zone. Sectigena sp. cf. S. crassa B GEYER 1990. Le I-X 234: Sectigena Zone. Sectigena crassa GEYER 1990. Le I-13: Sectigena Zone. Sectigena sp. cf. S. crassa B G EYER 1990. Le I-X 233: Sectigena Zone. Berabichia stenometopa G EYER 1990, Sectigena sdzuyi G EYER 1990 (type locality and stratum), hyoliths. Le I-17: Sectigena Zone. Hebediscus lemdadensis GEYER 1988. Le I-22: Sectigena Zone. Chancelloriids, Hebediscus lemdadensis GEYER 1988 (type locality and stratum), Antatlasia gemmea GEYER 1990 (type locality and stratum), Sectigena crassa GEYER 1990, Anabarella sp. Le I-X 235: Sectigena Zone. Hebediscus lemdadensis G EYER 1988, Sectigena iyouensis GEYER 1990 (type locality and stratum). Le I-X 236: Sectigena Zone. Chancelloriids, Hebediscus lemdadensis GEYER 1988, Sectigena sdzuyi GEYER 1990, echinoderm ossicles.
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Fig. 22. Trilobite Berabichia vertumnia GEYER 1990. Paratype, from Issafen Formation, Lemdad Syncline, section Le I-1/ 10a, sample horizon X 237a. Scale bar equals 1 mm. Naturmuseum Senckenberg, Frankfurt am Main, SMF 52266. Le I-X 232: Sectigena Zone. Antatlasia gemmea G EYER 1990, Sectigena crassa GEYER 1990. Le I-X 237a: Sectigena Zone. Archaeocyathans, chancelloriids, Hebediscus? sp., Iyouella contracta GEYER & PALMER 1995 (type locality and stratum), Lemdadella n. sp., Berabichia vertumnia G EYER 1990 (type locality and stratum), Hyolithellus sp. cf. H. micans BILLINGS 1872, hyoliths. Le I-X 237, X 237b: Sectigena Zone. Archaeocyathans, Hebediscus lemdadensis GEYER 1988, Sectigena crassa GEYER 1990, Le I-X 238: Sectigena Zone. Archaeocyathans, Hebediscus lemdadensis GEYER 1988, olenelloid trilobite, Lemdadella n. sp., Sectigena sdzuyi GEYER 1990, Berabichia vertumnia GEYER 1990, Berabichia sp., Anabarella sp., Conotheca? sp., hyoliths, Actinotheca sp., „Coleoloides“ sp. Le I-X 238a: Sectigena Zone. Chancelloriids, Antatlasia gemmea G EYER 1990, Sectigena crassa G EYER 1990, Anabarella sp., hyoliths. Le I-X 238b: Sectigena Zone. Archaeocyathans, Berabichia vertumnia GEYER 1990, Hyolithellus sp. cf. H. micans BILLINGS 1872. Le I-X 231: Sectigena Zone. Hebediscus lemdadensis G EYER 1988, Sectigena crassa G EYER 1990 (type locality and stratum), Sectigena sp., Antatlasia gemmea G EYER 1990. Le I-X 230: Sectigena Zone. Obolellida gen. et sp. indet., Marsaisia? n. sp., Sectigena sp. cf. S. crassa G EYER 1990, Antatlasia gemmea GEYER 1990. Le I-II/2-3: Probably Antatlasia guttapluviae Zone. Archaeocyathans, Berabichia vertumnia GEYER 1990. Le I-II/1-2: Archaeocyathans, Renalcis sp. Ounein B: Archaeocyathan level of DEBRENNE & DEBRENNE (unpubl. and this volume) and D EBRENNE et al. (1992). Paranacyathus patulus DEBRENNE, DEBRENNE & FAURE-MURET 1992, Archaeopharetra sp., Chouberticyathus clatratus DEBRENNE 1964, Protopharetra sp. aff. P. polymorpha BORNEMANN . Lemdad Formation Le Ia-X 241: Probably Antatlasia guttapluviae Zone: Calcareous algae, Berabichia sp. B, brachiopods, burrows. The following samples are from the archaeocyathan bioherm complex A. Probably Antatlasia guttapluviae Zone. Le I-I/1/16: Upper horizon of archaeocyathan bioherm. Probably Antatlasia guttapluviae Zone. Archaeocyathans, chancelloriids, Lemdadella? n. sp., Berabichia vertumnia GEYER 1990, Anabarella sp., Hyolithellus sp. cf. H. micans BILLINGS 1872, Renalcis sp. Le I-I/1/15: Renalcis sp. Le I-I/1/12: Archaeocyathans, chancelloriids, Berabichia vertumnia G EYER 1990, Hyolithellus sp. cf. H. micans B ILLINGS
1872, Renalcis sp. Le I-I/1/10: Main horizon of archaeocyathan bioherm. Archaeocyathans, Berabichia vertumnia GEYER 1990, Hyolithellus sp. cf. H. micans BILLINGS 1872, hyoliths, Renalcis sp. Ounein A: Archaeocyathan level of DEBRENNE & DEBRENNE (unpublished and 1995) and DEBRENNE et al. (1992). Afiacyathus ouneinensis DEBRENNE, DEBRENNE & FAURE-MURET 1992, Mennericyathus asper DEBRENNE, DEBRENNE & FAURE-MURET 1992, Polystillicidocyathus erbosimilis DEBRENNE 1959, Paranacyathus patulus DEBRENNE, DEBRENNE & FAURE-MURET 1992, Archaeopharetra sp., Chouberticyathus clatratus DEBRENNE 1964, Protopharetra sp. aff. P. polymorpha BORNEMANN. Le I-I/2/12: Archaeocyathans, chancelloriids, Hyolithellus sp. cf. H. micans BILLINGS 1872, lingulids. Le I-I/2/10: Archaeocyathans, Epiphyton sp., Renalcis sp. Le I-I/2/6: Archaeocyathans, Epiphyton sp. Le I-I/1/8: Archaeocyathans, undeterminable trilobites, Renalcis sp., undeterminable calcareous algae. Le I-I/1/7: Archaeocyathans, obolellid brachiopod, Hyolithellus sp. cf. H. micans BILLINGS 1872, Girvanella sp. Le I-I/1/5: Girvanella sp. Le I-I/4/7: Archaeocyathans, Hyolithellus sp. cf. H. micans BILLINGS 1872, acrotretids, hyoliths, trilobites, burrows. Le I-I/4/5: Archaeocyathans, Hyolithellus sp. cf. H. micans BILLINGS 1872, hyoliths, Renalcis sp. Le I-X 227: Probably Antatlasia guttapluviae Zone: Lemdadella? sp., Berabichia sp. A, undet. arthropod fragment, Hyolithellus sp. cf. H. micans BILLINGS 1872. Le I-X 228: Probably Antatlasia guttapluviae Zone: cf. Neltneria jacqueti (NELTNER & POCTEY 1950), Bondonella typica HUPÉ 1953, Berabichia sp. A GEYER 1990, Hyolithellus sp. cf. H. micans BILLINGS 1872. Le I-I/3T: Probably Antatlasia guttapluviae Zone. Berabichia vertumnia G EYER 1990 and other, undeterminable trilobites, Renalcis sp. F 1: Berabichia sp. B. F2: Berabichia sp. B. F 3: Marsaisia n. sp., antatlasiid genus and species. F 4: Berabichia sp. B. F 7: Bigotinops sp. F 5: Bigotinops sp. F 17: Daguinaspis sp. cf. D. ambroggii HUPÉ & A BADIE 1950, Bigotinops sp., hyoliths. F 16: Bigotinops sp. F 14, F 14a: Bigotinops sp. F 6: Bigotinops sp. F 15, L 15: Bigotinops sp. F 8, L 12: Bigotinops sp., Hyolithellus sp. cf. H. micans BILLINGS 1872. F 8b: Bigotinops sp. F 24, L 11: Bigotinops sp. F 23, L 10: Lemdadella sp. indet. S 2: Vetella sarfatiae SCHMITT 1979 (type locality and stratum) L 8: Lemdadella n. sp. L 7: Lemdadella n. sp. F 9, L 5: Lemdadella n. sp., burrows. S 1: Madiganites lemdadensis SCHMITT 1979 (type locality and stratum). F 22, L 4: Lemdadella n. sp. F 10a, L 3: Lemdadella n. sp. F 11b, L 2b, F 11a, L 2a: Lemdadella spectabilis SDZUY 1978, Hyolithellus sp. F 13, L 1: Lemdadella spectabilis SDZUY 1978 (type locality and stratum). Le Ia. Auxiliary section (section A2 sensu SDZUY 1978) Lemdad Formation F 12e: Botsfordiid new genus and species, „Bigotina“ sp., Lemdadella n. sp., undeterminate pelecypod, Pruvostina? sp. F 12d: Archaeocyathans, Lemdadella sp., Pruvostina? sp. F 12c: Inarticulate brachiopods, Hyolithellus sp., branched traces. F 12b: Lemdadella n. sp., Hyolithellus sp., burrows.
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Fig. 23. Detailed stratigraphic correlation of important sections through the transition from the Tata Group into the Feijas intérnes Group in the Lemdad syncline. Numbers refer to numeric sample horizons. Modified from GEYER (1990a). and GEYER et al. (1995) For location of section see Fig. 14.
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Stop 5. Section Le II Setting This north-dipping section is located about 0.8 km east of Aït Youb (Aït Iyou), and about 250 m east of section Le I (Lambert coordinates 225.6/425.7; Fig. 11). The base of section Le IIA is arbitrarily chosen at the top of the Issafen Formation and starts at the base of a small cliff in the formation. Section Le IIA is about 190 meters in thickness and ends at a fault zone (several tens of meters wide). The section continues, now labeled ‘Le IIB,’ north of the fault zone. Its top is overlain by red continental Cretaceous sedimentary rocks with an angular unconformity. Biostratigraphical significance The section was primarily chosen for this trip to illustrate the detailed, local fossil sequence available in the Tatelt Formation that can be used for refined biostratigraphic information. This section and section Le IV (see below) are the only ones that provide a rather complete fossil record through the Tatelt Formation. As this lithostratigraphic interval includes roughly the entire Hupeolenus Zone, it is, with Le IV, the only section that permits a crude evaluation of the Hupeolenus Zone interval. This zone is, due to unfavorable facies conditions, often poorly fossiliferous and/or crosses the boundary with the Jbel Wawrmast Formation so that an important part of the zone may be ab-sent in other sections. Section Le II proves that the Hupeolenus Zone is distinct but may represent a quite narrow time slice. Fossil evidence from section Le IV corroborates this information. In addition, the common occurrence of Protolenus (Hupeolenus) termierelloides together with Sectigena iyouensis proves that a biostratigraphic gap between the Hupeolenus and Sectigena Zones does not exist. Lithology and depositional environment of section Le II Section Le II resembles Le XI on the eastern flank of the Lemdad syncline in that the Tatelt Formation is mainly made up of volcanoclastic sandstones, arkoses, and quartz arenites deposited in a shallow marine setting. The fossils are usually accumulated in shell hash layers. Parts of the section with intercalated shales and generally finer-grained and slightly more reworked sediments appear to be deposited in a slightly more distal setting compared to section Le XI. The top of the formation includes intertidal, lenticular, flaser-bedded sandstones which document a generalised unidirectional flow direction to the southeast.
The base of the Jbel Wawrmast Formation contains two major condensed, highly fossiliferous horizons within the lower 8 m: one in the Cephalopyge Zone the other in the Ornamentaspis frequens Zone. By comparison with the axial zone of the Souss basin, where the same zones include at least 420 m of rocks (Tazlaft syncline, see below), the sections of the Lemdad syncline are extremely condensed. Tatelt Formation (Fig. 24) A thin (ca. 13–18 m) sandstone unit with considerable volcaniclastic debris overlies the shaly–volcanoclastic Issafen Formation. This unit was earlier thought to represent the regressive top of a ‘Grand Cycle’ and was termed the „Asrir Formation“ (GEYER, 1990a, GEYER & LANDING, 1995). It consists of a lower plane-bedded interval and an upper tidalite interval with bidirectional trough crosssets and calcareous nodules in the foresets. This „Asrir“ succession is lithologically similar to, coeval with, and referred to the Tatelt Formation as defined in the central Anti-Atlas (LANDING et al. 2006). The Tazlaft Formation (or lower part of the „grès terminaux“) is absent in the Lemdad area, and the Tatelt rests directly on the Issafen Formation. In section Le II, the Tatelt Formation consists of fineto coarse-grained arenites that are occasionally arkoses with euhedral feldspar crystals. The base of the formation is a 5–10 cm-thick polymict conglomerate with pebbles reworked from local units (e.g., phosphorite; greenish, friable rhyolite; feldspathic sandstone) and clasts derived from the Pan-African orogen (vein quartz and welded red rhyolite). Several layers of shales and conglomerates are intercalated. However, the conglomerates and graded bedding are restricted to the lower half of the section with graded bedding. Trough cross-bedding is very common. The bedding thickness varies between one centimeter and half a meter. Intercalations of shale or siltstone and their interbedding with medium-grained sandstones are common. The color throughout the formation is dark green, but grades to bluish, grey or medium green. Tuffs are occasionally intercalated into the siliciclastic sedimentary rocks. Compared with section Le XI, the more frequent shale and silt layers, absence of red colors and dunes, and comparatively rare basal conglomerates with finer components as well as the reduced thickness of the formation indicate deposition in a more distal environment. As in Le XI, body fossils are rare, and occur in occasional calcareous layers in the upper part. This upper part of the section (m 13–16) shows len-
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Fig. 25. Cliff-forming sandstones with high amount of volcanogenic material of the Tatelt Formation, overlain by „soft“ fine-grained sandstones with intercalated carbonate horizons of the Jbel Wawrmast Formation. Lemdad Syncline, Section Le II.
ticular flaser-bedding in medium-grained sandstone. These lenses are slightly calcareous and contain sands of differing grain-sizes up to gravel lags. The lowangle trough foresets, preserved only in lenses, indicate a flow direction predominantly to the southwest. However, there is also evidence for some reverse sediment transport. The top of the unit with the lenticular flaser-bedding is a 70 cm-thick arkose. The unit yields two fossil lags with a fauna dominated by the brachiopod Trematobolus, whereas trilobites are subordinate in this association. The trilobites indicate the Hupeolenus Zone as indicated by the pre-sence of the index fossil of the zone (Protolenus (Hupeolenus) hupei GEYER 1990). Interestingly, poorly preserved fragments of Paradoxides occur in this unit and zone. They are probably the oldest known Paradoxides remains known on Earth, and have allowed a reevalution of the global concept of the Lower–Middle Cambrian boundary (e.g., GEYER 1990d, 2005b; GEYER & LANDING 2005). The basal parts of the Tatelt Formation include remains of Sectigena crassa. This index fossil of the Sectigena Zone gives evidence that a major hiatus between the Sectigena and the Hupeolenus Zone (and Issafen and Tatelt Formations) is not developed in the Lemdad Syncline (GEYER et al. 1995). Recovery of Protolenus (Hupeolenus) hupei in the uppermost Issafen Formation Fig. 24 (opposite). Detailed profile of the Tatelt through Jbel Wawrmast Formations at section Le IIA illustrating grainsize profile, color changes, sedimentologic features, paleoenvironmental interpretation, and bio- and lithostratigraphic sequence. Thickness (left margin) in meters from the base of the section (top of Issafen Formation). See Fig. 27 for legend. Mainly based on HELDMAIER (1997) and modified from GEYER et al. (1995).
in this section not only shows that the range of Sectigena crassa extends into the Hupeolenus Zone (and thus into the lowest Middle Cambrian in the concept for Morocco), but also seems to suggest that a hiatus is not present between the formations. However, as detailed by LANDING et al. (2006), the Tatelt Formation directly overlies the Issafen Formation in the Lemdad Formation, without the intervening tidalite sandstones of the Tazlaft Formation as known in the Anti-Atlas. Thus, the thin polymict conglomerate at the base of the Tatelt Formation not only marks an abrupt change in facies from the deeper-marine conditions with HCS facies of the Issafen Formation to the tidalites of the lowest Tatelt Formation, but an important episode of uplift and unroofing of the Pan-African orogen that supplied vein quartz and welded rhyolites derived from the Pan-African orogen. This unconformity developed within the earliest Middle Cambrian Hupeolenus Chron. Geochronologic bracketing of the Hupeolenus Zone is presently undeveloped, but the complex history within this zone suggests a surprisingly long duration. The uppermost strata of the Tatelt Formation in Le II already belong to the Cephalopyge Zone so that the condensed formation nonetheless represent a considerable period in time but is not bound by major gaps.
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Jbel Wawrmast Formation (Fig. 24) The lower Jbel Wawrmast Formation (Brèche à Micmacca Member) includes a lower condensed arenaceous, phosphatic, brachiopod-trilobite limestone with corroded hardgrounds and upper phosphatic granule sandstone. A basal conglomerate (15 cm) with components up to 5 cm in diameter is developed locally, and is the local expression of the regional depositional sequence boundary that defines the base of the Jbel Wawrmast Formation (LANDING et al. 2006). The lower condensed interval of the Brèche à Micmacca Member is about 1.2 meter thick, nodular in habit and abounds with trilobite shells (a grainstone), as well as different reworked sediment particles. The middle of the horizon is formed by an intercalated 10 cm-thick, fine-grained arkose. This arkose is comparable to a black, 70 cm-thick fossiliferous arkose with masses of euhedral feldspars on top of the condensed horizon that underlies the yellowish green fine-grained sandstones of the typical Jbel Wawrmast facies. The base of the unit includes three fossiliferous grainstone horizons (up to 10 cm thick). A second condensed horizon lies about 2.5 m above the first condensed horizon. It also begins with a basal conglomerate (15 to 20 cm) with abundant glauconite grains and iron ooids and is similar to the second condensed horizon of section Le XI. The pack- to grainstones include reworked pebbles of the underlying yellowish green, fine-grained sandstones. In general, the Brèche à Micmacca Member (ca. 26 m) is identifiable as an interval characterized by several color changes from yellowish green to (lenticular) bluish green. These units are several meters thick and show a slightly increased carbonate content indicated by occasional, sometimes fossiliferous, nodular calcareous horizons. Burrows and other trace fossils (e. g., Teichichnus) occur sporadically. Body fossils are common within the lower 20 meters and are dominated by trilobites. Trilobites occur in thin shell lags or on single surfaces, and complete, often enrolled specimens are relatively common. Echinoderm ossicles are locally enriched in the lower several meters. Fossil shells are phosphatized from sample horizon Le II-81 upwards. The uppermost color change is seen at 127 m, where the top of the Brèche à Micmacca Member is drawn.
Fig. 26 (opposite). Detailed profile of the Jbel Wawrmast Formation at section Le IIB illustrating grain-size profile, color changes, sedimentologic features, paleoenvironmental interpretation, and bio- and lithostratigraphic sequence. Thickness (left margin) in meters from the base of the section (top of Issafen Formation). See Fig. 27 for legend. Mainly based on HELDMAIER (1997) and modified from GEYER et al. (1995).
At 136 m, a single thick shell bed with a slightly laminated 10 cm thick diagenetic underbed is found in the fine-grained sandstones. Nodular calcareous layers are rare higher up. Even when the marly nodules are fossiliferous, they reach large sizes (more than 0.5 meter width and 20 cm height). Mica flakes are common between 176 and 180 m. The mud- and siltstones of this interval are slightly laminated and include the trace fossil Teichichnus. A one meter thick massive bed with a 10 cm-thick shell layer in its center is found at the top of this unit. From 189 m upwards, the mud- and siltstones are again slightly laminated. The laminated top of section Le IIA likely correlates with the laminated unit just below the red-spotted ferruginous horizon (i.e., „Rotfleckenschiefer“ of BOUDDA 1968) of section Le XI. Fossil record Higher Jbel Wawrmast Formation Le IIA-141: Kymataspis arenosa Zone. Obolellid genus and species undeterminate, „Acrothele“ sp., Kymataspis arenosa GEYER 1990, Paradoxides (Acadoparadoxides) sp., ptychoparioid genus and species indeterminate, Conocoryphe (Parabailiella) sp. Le IIA-~80: Kymataspis arenosa Zone. Parasolenopleura n. sp. aff. P. conifrons W ESTERGÅRD 1953. Le IIA-58.75: Parasolenopleura lemdadensis GEYER 199. Le IIA-55: „Acrothele“ sp., Kymataspis sp. indet. Le II-1/45: Parasolenopleura lemdadensis GEYER 1998. Brèche à Micmacca Member Le IIA-29: Ornamentaspis frequens Zone. „Lingulella“ sp., „Acrothele“ sp., Condylopyge eli GEYER 1998, Ornamentaspis sp. cf. O. frequens GEYER 1990, „Paradoxides“ sp. indet., Agraulos sp. cf. A. arenosus SDZUY 1967. Le II-4T: Cephalopyge Zone. Acanthomicmacca sp. cf. A. neltneri HUPE 1953. Le II-4: Cephalopyge Zone. Trematobolus splendidus GEYER & MERGL 1995, Cobboldites sp. cf. C. tichkaensis GEYER 1988, Cephalopyge notabilis GEYER 1988, Latikingaspis sulcatus G EYER 1990, Latikingaspis alatus (HUPÉ 1953), Hamatolenus (Hamatolenus) marocanus (N ELTNER 1938), Latoucheia (Pseudolenus) ourikaensis (H UPÉ 1953), Paradoxides (Acadoparadoxides) nobilis GEYER 1998, Acanthomicmacca? sp., Acanthomicmacca sp. cf. A. neltneri HUPE 1953, Pelagiella sp. aff. P. lorenzi KOBAYASHI 1939, hyoliths, Marocella mira GEYER 1986, echinoderm ossicles. Le II-5: Cephalopyge Zone. Trematobolus splendidus GEYER & MERGL 1995, Cephalopyge notabilis GEYER 1988, Latikingaspis alatus (HUPÉ 1953), Latoucheia (Latoucheia) pusilla GEYER 1990, Paradoxides (Acadoparadoxides) cf. P. (A.) nobilis GEYER 1998, Acanthomicmacca sp. cf. A. neltneri HUPE 1953, Marocella mira GEYER 1986. Le II-7: Cephalopyge Zone. Trematobolus sp. cf. T. splendidus GEYER & MERGL 1995, Cephalopyge notabilis GEYER 1988. Le II-8: Cephalopyge Zone. Trematobolus sp. cf. T. splendidus G EYER & M ERGL 1995, Latoucheia (Latoucheia) pusilla GEYER 1990, Protolenus (Hupeolenus) sp. A. Le II-9: Cephalopyge Zone. Latikingaspis alatus (HUPÉ 1953). Tatelt Formation Le II-X 255: Hupeolenus Zone to Cephalopyge Zone. Mixture of sample horizons including roughly Le II-11 to Le II-9. Trematobolus splendidus GEYER & MERGL 1995, Cobboldites sp. cf. C. tichkaensis GEYER 1988, Cephalopyge notabilis G EYER 1988, Kingaspidoides? sp. C, Latoucheia (Latou-
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Le
Le
cheia) pusilla G EYER 1990, Protolenus (Hupeolenus) sp. cf. P. (H.) termierelloides GEYER 1990, Paradoxides (Acadoparadoxides) nobilis GEYER 1998, Gigoutella atlasensis H UPÉ 1953, Acanthomicmacca sp. cf. A. neltneri H UPE 1953, Conocoryphe (Parabailiella) sp., Jincella? sp., Helcionella capula G EYER 1986, Helcionella oblonga C OB BOLD 1921, Helcionella sp. aff. H. subrugosa ( D ’ORBIGNY 1850)?, Leptostega? sp. aff. L. cingulata (COBBOLD 1921), Pelagiella sp. aff. P. lorenzi K OBAYASHI 1939, hyoliths, Marocella mira GEYER 1986. II-11: Cephalopyge Zone. Trematobolus splendidus GEYER & MERGL 1995, Saesorthis sp. cf. S. simplicissima (MERGL 1982), Kingaspidoides obliquoculatus GEYER 1990, Kingaspidoides? sp. C, Protolenus (Hupeolenus) sp. cf. P. (H.) hupei GEYER 1990, Protolenus (Protolenus) densigranulatus GEYER 1990, Latoucheia (Latoucheia) pusilla GEYER 1990, Acanthomicmacca? sp., Acanthomicmacca sp. cf. A. neltneri HUPE 1953, Paradoxides (Acadoparadoxides) nobilis GEYER 1998, hyoliths. II-11u: Probable Cephalopyge Zone. Latoucheia (Latoucheia) pusilla GEYER 1990, Protolenus (Protolenus) densigranulatus GEYER 1990, Protolenus (Protolenus) densigranulatus GEYER 1990, Kingaspidoides obliquoculatus GEYER 1990, Paradoxides (Acadoparadoxides) nobilis G EYER 1998. II-12/11: Probably Cephalopyge Zone. Protolenus (Protolenus) densigranulatus GEYER 1990. II-12/12: Hupeolenus Zone or probably Cephalopyge Zone. Sponge spicules, undeterminable obolellids, Calodiscus n. sp. A, Tschernyshevioides? sp., Kingaspidoides obliquoculatus G EYER 1990, Latoucheia (Latoucheia) pusilla G EYER 1990, Protolenus (Protolenus) densigranulatus G EYER 1990, Paradoxides (Acadoparadoxides) nobilis GEYER 1998, Latouchella sp. cf. L. iacobinica GEYER 1986, „Helcionella“ sp. indet. II-12/15: Hupeolenus Zone. Trematobolus splendidus GEYER & MERGL 1995, Protolenus (Hupeolenus) termierelloides GEYER 1990. II-12/21: Hupeolenus Zone. Trematobolus sp. cf. T. splendidus GEYER & MERGL 1995, Protolenus (Protolenus) densigranulatus GEYER 1990, Protolenus (Hupeolenus) hupei GEYER 1990, Latoucheia (Latoucheia) pusilla GEYER 1990, Kingaspidoides obliquoculatus GEYER 1990, Kingaspidoides? sp. C, „Paradoxides“ sp. II-12/25: Hupeolenus Zone. Trematobolus? sp., Sectigena crassa GEYER 1990, Protolenus (Hupeolenus) termierelloides GEYER 1990 (type locality and stratum), Protolenus (Hupeolenus) hupei G EYER 1990 (type locality and stratum). II-12/300: Hupeolenus Zone. Trematobolus sp., Kingaspidoides obliquoculatus GEYER 1990, Protolenus (Hupeolenus) hupei G EYER 1990, Protolenus (Hupeolenus?) sp., Protolenus (Protolenus) densigranulatus GEYER 1990 (type locality and stratum), „Paradoxides“ sp., Latouchella sp., Marocella mira GEYER 1986.
Le II-12/400: Hupeolenus Zone. Trematobolus sp., undeterminable olenelloid trilobite, Protolenus (Hupeolenus) hupei GEYER 1990, hyoliths. Le II-12/03: Sectigena Zone. Alisina? sp., Sectigena crassa GEYER 1990.
Section Le IIB (Fig. 26) Section Le IIB includes a lateral continuation of the Jbel Wawrmast Formation from section Le IIA. It begins north of a fault zone of about 50 meter in width, which intersects the upper part of section Le II. The amount of displacement is unclear. The lithology can tentatively be correlated with the upper part of section Le XI. If this lithologic correlation is correct, approximately 125 m of sandstone present in section Le XI are lacking in Le II (including the red-spotted ferruginous horizon termed „Rotfleckenschiefer“ by BOUDDA 1968). The Jbel Wawrmast Formation consists of green finegrained sandstones which were deposited on a distal shelf. The sedimentary rocks are more-or-less laminated and thin-bedded in the lower part (85 m) of the section. Lamination and bedding are coupled with changes in the mica content. The more mica present, the thicker are the individual beds. This is very similar to section Le XI, but precise correlation of these „bedding units“ between the sections is not possible. Extremely large and thick calcareous nodules are occasionally present, as well as color changes from the usual yellowish to bluish green. With the disappearance of abundant mica from 85 m to the top (152.5 m) of the section, more sedimentary features are observable. Burrow-churning (with the presence of Teichichnus and large horizontal traces) becomes particularly common. Trilobites („Paradoxides“) and echinoderms (eocrinoids) are found either as scattered single specimens or are disarticulated in shell hash beds. Isolated trilobite specimens occur in horizons with relatively abundant fossils which sometimes reach several meters in thickness. At 88.95 m, a roughly one centimeter thick volcanic ash is preserved. From 120 m upward bluish green, slightly calcareous layers occur again.
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Stop 6. Section Le IV The short, north-dipping Le IV section (Fig. 23) includes a succession from the Issafen Formation to the lower Jbel Wawrmast Formation. Of particular interest is the transition from the Issafen to the Tatelt Formation (i.e., „Asrir Formation“ in GEYER et al. 1995), which is well exposed. The section is particularly important for the biostratigraphy of the Issafen–Tatelt contact. HCS sandstone beds in the Issafen Formation yield numerous trilobites,
particularly specimens of Antatlasia, Berabichia, and Sectigena. Bed Le IV-9 shows an overlap of the ranges of Sectigena and Hupeolenus, and Protolenus (Hupeolenus) termierelloides has been found together with Sectigena iyouensis on the same bedding plane at the base of the section. GEYER et al. (1995) noted that this persistence of Sectigena into the Hupeolenus Zone proves that a stratigraphical gap with a possible lack of one or more zones is not present at the unconformity
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Fig. 27. Legend for Figs. 24, 26, 28, 30, 33, and 38.
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marked by the thin polymict conglomerate at the base of the Tatelt Formation. However, the absence of the Tazlaft Formation at all sections in the Lemdad Syncline and evidence for epeirogenically driven tilting that led to the differential erosion of 2.5 m of the upper Issafen Formation between section Le II and Le IV provide evidence of a major unconformity and a relatively long hiatus within the Hupeolenus Zone (LANDING et al. 2006, Fig, 28). The implications of this interpretation are 1) that the Hupeolenus Zone had a long duration and 2) that Cambrian trilobite zones did not succeed each other with a „clock-like“ regularity, and some, as the Hupeolenus Zone, represent times of relative stasis. Fossil record Upper Jbel Wawrmast Formation Le IV-1/10: Kymataspis arenosa Zone. Kymataspis arenosa GEYER 1990, ptychopariid new genus and species D. Le IV-1/8: Kymataspis arenosa Zone. Kymataspis arenosa GEYER 1990. Le IV-1/7: Kymataspis arenosa Zone. Kymataspis arenosa GEYER 1990, Paradoxides (Acadoparadoxides) sp., echinoderm ossicles. Brèche à Micmacca Member Le IV-2: Ornamentaspis frequens Zone. Paradoxides (Acadoparadoxides) sp., Ornamentaspis frequens GEYER 1990,
Parasolenopleura? sp., ptychopariid new genus and species, „Syspacephalus“ sp. Le IV-3: Probable Cephalopyge Zone. Trematobolus sp. cf. T. splendidus GEYER & MERGL 1995, trilobite fragments. Le IV-4: Probable Cephalopyge Zone. Chancelloriids, Trematobolus sp. cf. T. splendidus GEYER & M ERGL 1995, trilobite hash. Le IV-8: Probable Cephalopyge Zone. Trematobolus sp. cf. T. splendidus GEYER & MERGL 1995, trilobite hash. Tatelt Formation Le IV-9: Probable Hupeolenus Zone. Trematobolus sp. cf. T. splendidus GEYER & MERGL 1995, Protolenus (Hupeolenus) sp. cf. P. (H.) termierelloides GEYER 1990. Le IV-9B: Hupeolenus Zone. Protolenus (Hupeolenus) termierelloides GEYER 1990, Sectigena iyouensis GEYER 1990, Sectigena sdzuyi GEYER 1990. Issafen Formation Le IV-11: Sectigena Zone. Hebediscus lemdadensis GEYER 1988, Sectigena iyouensis GEYER 1990. Le IV-15: Sectigena Zone, Sectigena crassa GEYER 1990, Antatlasia gemmea GEYER 1990. Le IV-17: Sectigena Zone. Hebediscus lemdadensis GEYER 1988, Antatlasia gemmea GEYER 1990, Sectigena crassa GEYER 1990. Le IV-19: Sectigena Zone. Antatlasia gemmea GEYER 1990, Sectigena crassa GEYER 1990. Le IV-22T: Antatlasia guttapluviae Zone or Sectigena Zone. Berabichia stenometopa G EYER 1990. Le IV-23: Antatlasia guttapluviae Zone or Sectigena Zone. Antatlasia gemmea G EYER 1990. Le IV-27T: Probable Antatlasia guttapluviae Zone or Sectigena Zone. Berabichia stenometopa GEYER 1990.
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Stop 7. Section Le XI Setting This southwest-dipping section is located about 2 km south of Taliwine n’ Aït-Al Mimoun in the upper Lemdad valley (Lambert coordinates 235/425,3) (Figs. 28–30). Previous studies dealt with the archaeocyathans (DEBRENNE & DEBRENNE 1975; DEBRENNE et al. 1992), which are the best studied in the region and crucial for the biostratigraphical assignment of the younger archaeocyathan faunas of Morocco (bioherme de la cascade; DEBRENNE & DEBRENNE 1975, 1995). The base of the measured part of the section (Fig. 30) is in the Lemdad Formation at the rim of the large travertine mass that forms the hill crest to the northeast of the section. The Cambrian rocks are overlain by red, Cretaceous sediments with an angular unconformity at 540 m. Lithologic sequence Summary Section Le XI is a reference section for the eastern flank of the syncline and has particularly good exposures through the transition from the Tatelt into the Jbel Wawrmast Formation.
Deposits assigned to the Tatelt Formation directly overlie an archaeocyathan bioherm of the lowest Issafen Formation, and which rests on the Lemdad Formation. The Issafen Formation, known in other parts of the Lemdad syncline, is almost entirely eroded away and is represented only by the bioherm. The Lemdad Formation is dominated by volcaniclastic-rich arkoses that were deposited in shallow, wave-dominated environments. All of the material was derived from a dominantly acidic volcanic complex located close to the locality. A 10 cm-thick greenish, rhyolitic ash 10.8 m above the base of the section has provided a U–Pb zircon age of 517 ± 1.5 Ma in the upper Antlasia guttapluviae Zone (LANDING et al. 1998). This complex is overlain by deposits which were strongly influenced by tidal currents. The conspicuous archeocyathan bioherm (bioherme de la cascade), representing the Issafen Formation, has become the name-giver for the section (DEBRENNE & DEBRENNE 1975; BOUDDA et al. 1979). The archaeocyathans, stromatolitic limestone clasts, flat pebble conglomerates, an oolite bed, and intercalated purple shales, suggest an extremely shallow, nearshore marine depositional environment with variable water energies. Intercalated ash layers in the immediately overlying siliciclastic units
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indicate continuing volcanic activity during deposition of the Tatelt Formation. The Brèche à Micmacca Member of the Jbel Wawrmast Formation includes two conspicuous fossiliferous limestones in the lower part that represent condensed horizons. The biostratigraphical data of these lower 8 meters indicate the Cephalopyge notabilis to lower Ornamentaspis frequens Zones. As in the other sections in the Lemdad syncline, this interval is generally extremely condensed in comparison with those in the central Souss Basin (e. g., the Tazlaft syncline where more than 420 m of fine-grained sandstone were formed). The upper boundary of the Jbel Wawrmast Formation is difficult to recognize, and stratigraphic equivalents of the Jbel Afraou Formation occur as a Jbel Wawrmast facies in the section. The Bailiella Formation, which also generally resembles the Jbel Wawrmast Formation, cannot be distinguished with certainty. However, Bailiella levyi, characteristic of the formation in the central Anti-Atlas, is found in the upper part of the section. Lemdad, Issafen and Tatelt Formations The Tatelt Formation in section Le XI has a thickness that is roughly the same as that in the western flank of the Lemdad syncline. The Tatelt Formation in section Le XI overlies a thick „volcano-detrital complex,“ which correlates with the upper part of the Lemdad Formation in other parts of the Lemdad syncline, and the bioherme de la cascade (DEBRENNE & DEBRENNE 1975). The bioherme de la cascade intervenes between the Lemdad and Tatelt Formations, and represents the Issafen Formation. All higher parts of the Issafen Formation are eroded away (Fig. 28). This bioherm has been interpreted to contain the youngest archaeocyathans of the region (DEBRENNE & DEBRENNE 1975; DEBRENNE et al. 1992) and has been regarded as crucial for the biostratigraphy of the youngest archaeocyathan faunas of Morocco. However, this correlation has been rejected (GEYER et al. 1995) based on the presence of Berabichia sp. cf. B. vertumnia in the bioherm. In addition, near section Le X (short distance south of section Le XI), where the archaeocyathan limestones is transitional into shales, are overlain by trilobite-bearing calcareous and sandy intercalations referable to the Sectigena Zone. These data indicate that this archaeocyathan bioherm (Ounein C) is not distinctly younger than those described from section Le I and is most probably referable to the Antatlasia guttapluviae Zone. This lower part of section Le XI consists of quartz and volcanoclastic arenites and arkoses. The deposits include reworked volcanic material, which is characterized by abundant, almost euhedral feldspars. Larger rhyolitic clasts are found occasionally. Euhedral feldspars
may compose some beds almost exclusively. The thickness of the beds varies from several centimeters up to 2 meters. The grain size varies from fine silt to coarse sand. The top of the Lemdad Formation includes conglomerate layers. The sandstones are usually green, but the silty particles are reddish. Single beds are often normally graded. Occasionally, the coarser base of a bed consists of a conglomerate of reworked pebbles of the underlying sediments mixed with rhyolite fragments. Lamination, trough cross-bedding, ripples, and, rarely, HCS occur. Volcanic ash layers of about 1 cm thickness, sometimes even thinner than 1 mm, are intercalated among the tuffitic or „volcanodetritic“ material. This sedimentary succession of the Lemdad Formation is typical for a tidally dominated coast (trough and planar cross bedding, coarse sands, parallel lamination and mostly poorly sorting) with channels (coarse, pebbly and erosional base) and with an important component of wave activity. Megaripples and dunes are found from 24 m upward. The archeocyathan bioherm of the lowest Issafen Formation grew on a basal conglomerate with flat, slightly rounded, sandstone cobbles and boulders (10 to 30 cm in length) and rhyolite fragments. This setting is similar to that of the upper archeocyathan bioherm in section Le I. The upper surface of the archaeocyathan bioherm has a strong relief, which is filled by fine-grained, partly laminated sands of the unconformably overlying Tatelt Formation. These are overlain by coarse sandstones with calcareous nodules that show megaripples, large troughs, and dunes. Intercalated conglomerates have limestone clasts. Several layers of purple siltstones are intercalated. The top of the Tatelt Formation is formed by a roughly one meter-thick, ferruginous oolite. These upper beds of the Tatelt Formation appear to have been deposited in a very shallow, high-energy environment. Megaripples (possibly formed on a flood ramp) and ebb channels with trough cross-stratification and the occasional reworking of barely lithified sediments (intraclasts) indicate a strong tidal influence. Jbel Wawrmast Formation The lower boundary of the Brèche à Micmacca Member of the Jbel Wawrmast Formation is a coarse-grained, reddish, fossiliferous sandstone which is intercalated between two 6 cm-thick purple shale layers. The lower of these shale layers has an ash bed about 1 cm-thick. The coarse sandstone consists of reworked sand-grains of the underlying beds and contains reworked clasts of stromatolitic crusts. A lower, ca. 1.0 m-thick, condensed limestone horizon that sits on top of the sandstone is conspicuous. This bed contains abundant shell frag-
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Fig. 28. Correlation of the Lower--Middle Cambrian interval in the Lemdad sections Le I, Le II, and Le XI, illustrating the differential erosion of the Issafen Formation. Further explanations in the text. From LANDING et al. (2006, fig. 7), with minor modifications.
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ments, stromatolitic crusts, and occasional shale intercalations. The fauna includes numerous brachiopods and echinoderm ossicles, but trilobites typical of the Cephalopyge notabilis Zone are relatively rare and have low diversity. The limestone becomes nodular when the shale intercalations increase in number. Intensive burrowing increases upward with the grain-size of the matrix, and sand-size particles clearly increase towards the top of the limestone. The second limestone horizon has a diverse fauna with trilobites of the Ornamentaspis frequens Zone. Chancelloriids, acrotretid brachiopods, helcionellids and other small shelly fossils are frequent. Echinoderm ossicles are extremely abundant in parts of the bed. The fauna adds to the largest diversity of Cambrian fossils known from Morocco. However, the horizon is clearly amalgamated and consists of at least four different phases of shell and sclerite accumulations. Small shelly fossils and trilobite sclerites are abraded in different ways so that a condensation is indicated. However, biostratigraphically useful fossils all indicate the Ornamentaspis frequens Zone so that serious stratigraphic problems because of condensation and amalgamation (as discussed by ÁLVARO & CLAUSEN 2004, 2006) does not exist in the Lemdad sections. A typical succession, including the color changes that are characteristic of the Brèche à Micmacca Member, condensed horizons, and fossiliferous carbonate nodules, occurs up to 80 m. In addition, volcanic ashes are also preserved in the Brèche à Micmacca Member. The trilobites record an extremely condensed section for the lower 8 meters (up to the second condensed horizon) compared with the center of the Souss Basin (e. g., the Tazlaft Syncline). Thus, the lithology of the Brèche à Micmacca Member is clearly identifiable. By contrast, the upper boundary of the Jbel Wawrmast Formation is not determinable because the facies of the Jbel Afraou Formation are not developed as in the southwestern High Atlas. Elsewhere, the base of the Jbel Afraou Formation is defined by the lowest occurrence of HCS-bearing sandstones, and this facies is not developed in the Lemdad syncline. A characteristic Jbel Wawrmast Formation facies continues upward into strata correlative with the Jbel Afraou Formation in the central and eastern Anti-Atlas. Upper part of the section (Figs. 29, 30) A slight facies change, featuring increased lamination coupled with an increased carbonate content, locally expressed by development of large carbonate nodules or nodular layers, as well as a bluish color occurs 26 m below the „red-spotted shale“ („Rotfleckenschiefer“, see below). The mica content increases as well in these layers.
Fig. 29. Section Le XI as seen from the hill crest on the travertine mass at the base of the sequence. Numbers refer to the measurements given in Fig. 30. Marker and fossil horizons are noted (RFS = „Rotfleckenschiefer“). Further explanations are in the text.
The only well developed lithostratigraphic marker horizon in the higher part of the sequence is found at 217 m. It is a fossiliferous, yellowish to red fine-grained sandstone (quartz arenite to arkose, up to one meter thick) with characteristic iron nodules. This bed, cemented by dolomite, is erosionally resistant and topographically prominent. It has been called „Rotfleckenschiefer“ („redspotted shale“; Fig. 29) by BOUDDA (1968), although it is lithologically a sandstone. On its top is a nonfossiliferous sandstone with a thickness of 15 to 20 cm, also cemented by dolomite. Primary sedimentologic features are not present. Although this horizon is easily recognizable at the eastern flank of the Lemdad syncline, it is not yet correlatable into other regions. The mica content increases generally above the „redspotted shale.“ The amount of mica is reflected in the decreased bedding thickness which lies between one and two millimeters. Horizontal burrows of up to 2 cm in diameter occur sporadically. Several syndepositional mass movements are displayed between 262 and 282 m (Figs. 30). Slide scarps
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91
92
93
Fig. 30. Detailed sequence of section Le XI illustrating grain-size profile, color changes, sedimentologic features, paleoenvironmental interpretation, and bio- and lithostratigraphic succession. Bio- and lithostratigraphic boundaries of the upper part of the section in are explained in the text and on Fig. 29. Thickness (left margin) in meters from the base of the section. See Fig. 27 for legend. Based mainly on HELDMAIER (1997) and modified from GEYER et al. (1995).
with oblique discontinuity surfaces and a relief of up to 1.1 meter, with overlying slightly coarser and intensively laminated sandstones, indicate extensive mass movements down slope. At 302.5 m the rock color changes from green into grey. Phosphatized, reworked trilobites, as well as little glauconite balls, occur. The sandstone is cemented by dolomite. On top of this unit is a second condensed sandstone horizon, a 20 to 50 cm-thick shell hash bed. The enrichment of glauconite is noteworthy. About 5.5 meters of green, fine-grained sandstones are present above this shell hash, and this interval has many small glauconite balls (1 to 3 mm in diameter). Yellowish, green fine-grained sandstones reappear higher in thre interval. A third condensed sandstone horizon is present at 363 m. It is a 47 cm-thick calcareous, fine-grained sandstone that frequently bears phosphatized trilobite shells. The color of the overlying sandstones changes repeatedly from yellowish green to bluish green. Isolated, often very large carbonate nodules are found higher. A fourth condensed sandstone horizon is located at 424 m. The basal 10 cm are made up of a trilobite–echinoderm grainstone. The upper 17 cm are dominated by sandy layers enriched with glauconite grains. The inter-
val between 425 and 450 m contains non-laminated, finegrained, green sandstones. Burrowing increases slightly in this interval. Teichichnus is abundant at 444 m. In addition, trilobites from this interval indicate one of the few certain records of the Pardailhania Zone in Morocco (with Pardailhania hispanica SDZUY 1958 present between 430 and 440 m). From 451.6 m upward, the sandstones are laminated. Fossiliferous, calcareous fine-grained sandstone layers occur sporadically. Carbonate nodules are absent. Two 40 cm-thick arkoses with euhedral feldspars and mediumand coarse-grained sandstones, cemented by carbonate, are intercalated at 532 m. An angular unconformity with overlying, red, Cretaceous rocks truncates the section at 540 m. Fossil record Equivalents of Bailiella Formation X 222: Strata with Bailiella levyi. „Jamesella“ sp. A, „Paradoxides“ sp., Conocoryphe (Parabailiella) sp., Kingaspidoides sp., Conocoryphe sp. cf. C. pseudooculata MIQUEL 1905, Sucocystis quadricornuta F RIEDRICH 1993 (type locality and type stratum), additional isolated echinoderm ossicles. X 207a, X 70: Strata with Bailiella levyi. „Paradoxides“ sp., Bailiella sp. cf. B. levyi (MUNIER-CHALMAS & BERGERON 1889), Bailiaspis sp.
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Upper part of Jbel Wawrmast Formation Le XI-439: Pardailhania Zone. Ornamentaspis? sp. aff. O. destombesi G EYER 1990, „Paradoxides“ sp., ptychpariid genus and species uncertain, Conocoryphe (P.) sp. cf. C. (P.) sebarensis SDZUY 1967, Pardailhania hispanica SDZUY 1958. Le XI-431: Probable Pardailhania Zone. Paradoxides (Eccaparadoxides?) sp. Le XI-408,5: Acrotretid genus and species undeterminate, „Acrothele“ sp., Kingaspidoides? sp. undet., echinoderm ossicles. X 209, X 55: Badulesia Zone. Badulesia tenera Zone. Kymataspis arenosa GEYER 1990, „Paradoxides“ sp., Conocoryphe (Parabailiella) sp. aff. C. (P.) schmidti SDZUY 1957, Badulesia tenera (HARTT 1868), Ctenocephalus (Harttella) sp. aff. C. (H.) antiquus T HORAL 1946. X 208: Badulesia Zone. „Paradoxides“ sp., Parasolenopleura sp. aff. P. aculeata (ANGELIN 1851), Conocoryphe (Parabailiella) sp. aff. C. (P.) schmidti SDZUY 1957, Badulesia tenera (HARTT 1868), echinoderm ossicles. Le XI-145: Badulesia Zone. Badulesia granieri (H ARTT 1868) (det. GOZALO & LIÑAN). Le XI-138: Badulesia Zone. Paradoxides (Eccaparadoxides) asturianus SDZUY 1967, Conocoryphe (Parabailiella) sp. aff. C. (P.) schmidti SDZUY 1957, Badulesia tenera (HARTT 1868) (det. GOZALO & LIÑÁN). Le XI-136: Badulesia Zone. Conocoryphe (Parabailiella) sp. aff. C. (P.) schmidti SDZUY 1957, Badulesia tenera (HARTT 1868). Le XI-135: Badulesia Zone. Badulesia tenera (H ARTT 1868) (det. GOZALO & LIÑAN). Le XI: 108: Parasolenopleura lemdadensis GEYER 1998. Le XI-107.5: Parasolenopleura lemdadensis GEYER 1998. Le XI-103: Parasolenopleura lemdadensis GEYER 1998 Le XI-83: Parasolenopleura lemdadensis GEYER 1998. Le XI-1/57: Parasolenopleura lemdadensis GEYER 1998. Le XI-1/55: Probable Kymataspis arenosa Zone. Kymataspis arenosa GEYER 1990, Parasolenopleura lemdadensis GEYER 1998. Le XI-1/54: Probable Kymataspis arenosa Zone. Kymataspis arenosa GEYER 1990. Le XI-100: Kymataspis Zone. Paradoxides (Eccaparadoxides) asturianus SDZUY 1967 (det. GOZALO & LIÑÁN), Badulesia? sp. A. Le XI-98: Kymataspis Zone. Paradoxides (Eccaparadoxides) asturianus SDZUY 1967 (det. GOZALO & LINÁN). Le XI-1/30-40: Kymataspis arenosa Zone. Kymataspis arenosa GEYER 1990, Paradoxides (Acadoparadoxides) sp. aff. A. sacheri B ARRANDE 1854, Parasolenopleura lemdadensis GEYER 1998. Le XI-1/25: Kymataspis arenosa Zone. Parasolenopleura? sp. Jbel Wawrmast Formation, Brèche à Micmacca Member Le XI-25/50: Ornamentaspis frequens Zone. Ornamentaspis sp. cf. O. angustigena GEYER 1990. Le XI-24e: Ornamentaspis frequens Zone. Ptychopariid new genus and species, Parasolenopleura? sp. Le XI-X 214, Le XI-51.5: Ornamentaspis frequens Zone. Chancelloria maroccana, acrotretids, antagmid genus and species A, Cambrophatictor cataractis G EYER 1998, Conocoryphe (Parabailiella) sp., Atopiaspis tikasraynensis G EYER 1998, Ornamentaspis crassilimbata G EYER 1990, Kingaspidoides sp. aff. K. laetus G EYER 1990, „Paradoxides“ sp., Bailiella n. sp. A, Bailiella n. sp. B, Latouchella sp. A, Latouchella sp. cf. L. iacobinica GEYER 1986, Latouchella comma G EYER 1986 (type locality and stratum), Bemella sp., Stenothecopsis? n. sp., new helcionellid genus and species E, new helcionellid genus and species G, Pelagiella atlasensis GEYER 1986 (type locality and stratum), Pelagiella sp. A, hyoliths, „Actinotheca“ sp., Hyolithellus sp., various echinoderm ossicles, problematic spherical SSFs.
Le XI-24c: Ornamentaspis frequens Zone. Chancelloria maroccana, acrotretids, Trematobolus splendidus GEYER & MERGL 1995, Ornamentaspis crassilimbata GEYER 1990, Kingaspidoides sp. aff. K. laetus GEYER 1990 ?, „Paradoxides“ sp., „Syspacephalus“ sp., Cambrophatictor cataractis GEYER 1998, Bailiella dilatata G EYER 1998, Parasolenopleura? sp., Pelagiella atlasen-sis GEYER 1990, hyoliths, echinoderm ossicles. Le XI-24b: Ornamentaspis frequens Zone. Trematobolus sp., Chancelloria maroccana, acrotretids, Ornamentaspis crassilimbata GEYER 1990, Parasolenopleura? sp., Bailiella? n. sp. A, hyoliths, echinoderm ossicles. Le XI-24a: Ornamentaspis frequens Zone. Trematobolus splendidus GEYER & M ERGL 1995, Ornamentaspis crassilimbata GEYER 1990, „Syspacephalus“ sp., Parasolenopleura? sp. Le XI-51,9. Chancelloria maroccana, acrotretids, Ornamentaspis crassilimbata G EYER 1990, ptychopariid trilobite, Latouchella comma GEYER 1986, Yochelcionella sp., Pelagiella atlasensis GEYER 1986, hyoliths, Hyolithellus sp., echinoderm ossicles, problematic ESFs. Le XI-51,65. Chancelloria maroccana, acrotretids, Ornamentaspis sp., ptychopariid trilobite, hyoliths, echinoderm ossicles. Le XI-23/71: Ornamentaspis frequens Zone. Acrothele sp., Ornamentaspis crassilimbata GEYER 1990, Kingaspidoides sp. aff. K. laetus GEYER 1990. Le XI-23/61: Ornamentaspis frequens Zone. Ornamentaspis frequens Zone. Ornamentaspis crassilimbata G EYER 1990, Kingaspidoides sp. aff. K. laetus GEYER 1990, paradoxidid genus and species uncert. Le XI-23/58: Ornamentaspis frequens Zone. „Kingaspis“ sp. indet. Le XI-23/55: Ornamentaspis frequens Zone. „Lingulella“ sp., „Acrothele“ sp., acrotretids, Ornamentaspis crassilimbata GEYER 1990, Kingaspidoides sp. aff. K. laetus GEYER 1990, Protoleninae gen. et sp. incert., Bailiella? sp. L XI-X 219: Ornamentaspis frequens Zone. „Acrothele“ sp., ptychopariid new genus and species, Ornamentaspis crassilimbata GEYER 1990 (type locality and stratum), Kingaspidoides sp. aff. K. laetus GEYER 1990, „Paradoxides“ sp., Parasolenopleura n. sp., Marocella mira GEYER 1986. Le XI-23a: Probable Ornamentaspis frequens Zone. Trematobolus sp. cf. T. splendidus GEYER & MERGL 1995, „Acrothele“ sp., Kingaspidoides sp. aff. K. laetus GEYER 1990, „Paradoxides“ sp. Le XI-42.7: Probable Cephalopyge Zone. Chancelloria maroccana, Pelagiella atlasensis GEYER 1990, echinoderm ossicles. Le XI-41,5: Cephalopyge Zone. Trematobolus splendidus GEYER & MERGL 1995, trilobite sclerites, Hyolithellus sp., echinoderm ossicles. Le XI-22b: Cephalopyge Zone. Trematobolus splendidus GEYER & M ERGL 1995, Pseudocobboldia pulchra (H UPÉ 1953), „Paradoxides“ sp., Vallatotheca n. sp., Anabarella sp., hyoliths, Marocella mira GEYER 1986, echinoderm hash. Le XI-X 213: Cephalopyge Zone. Trematobolus splendidus GEYER & MERGL 1995, Saesorthis sp. cf. S. simplicissima (MERGL 1982), Acanthomicmacca sp. cf. A. neltneri H UPE 1953, „Paradoxides“ sp., hyoliths, Marocella mira GEYER 1986, echinoderm ossicles. Le XI-21/60: Probably Cephalopyge Zone. Trematobolus splendidus GEYER & MERGL 1995, Marocella mira GEYER 1986. Le XI-21/55: Probably Cephalopyge Zone. Trematobolus splendidus GEYER & MERGL 1995, Protolenus (Protolenus) densigranulatus GEYER 1990. Le XI-21/40: Probable Cephalopyge Zone. Protolenus (Hupeolenus) termierelloides GEYER 1990. Tatelt Formation Le XI-39.8: Probable Hupeolenus Zone. Trematobolus sp. cf. T. splendidus GEYER & M ERGL 1995, Marocella mira GEYER 1986, Protolenus (Hupeolenus) sp. indet. Le XI-X 212, Le XI-37.6: Hupeolenus Zone. Trematobolus sp.
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cf. T. splendidus GEYER & M ERGL 1995, Marocella mira GEYER 1986, Protolenus (Hupeolenus) hupei GEYER 1990, Protolenus (Hupeolenus) termierelloides GEYER 1990. Issafen Formation Ounein C: Archaeocyath bioherm „La Cascade“ by DEBRENNE & DEBRENNE (1975, 1995) and DEBRENNE et al. (1992). Afia-
cyathus alloiteaui DEBRENNE 1964, Mennericyathus asper DEBRENNE, DEBRENNE & FAURE-MURET 1992, Erismacoscinus calathus (BORNEMANN), „Syringocoscinus,“ „Antoniocoscinus“, „Churanicyathus“, Dictyocyathus verticillus (BORNEMANN). Le XI-31: Most probably Antatlasia guttapluviae Zone. Berabichia sp. cf. B. vertumnia GEYER 1990.
December 4
Stop 8. BOU IGHIR Road cuts on the western side of the road from Taroudant to Igherm close to the village of Bou Ighir expose the upper part of the Adoudou Formation and illustrate the parasquences of the upper most Tifnout Member. This section is very similar to that in the Sdas valley (slightly to the west) which was studied (particularly its isotopic signatures) in detail by LATHAM (1990), LATHAM & RIDING (1990), and MALOOF et al. (2005). The Sdas valley succession serves as a standard for the earliest Cambrian carbon isotope curve of Morocco and allows correlation into other Cambrian continents (see MAGARITZ
et al., 1991, BRASIER 1991, MALOOF et al. 2005, and discussion in the introductory part (GEYER & LANDING, this volume). The roadcuts show carbonate-dominated, shallowingupward, meter-scale parasequences. They generally begin with variably colored siltstones and marlstones, which tend to be very thin at the Bou Ighir locality, followed by laminated micritic limestones and dolostones. Locally, laminated carbonates with microbial mats form the tops of the parasequences.
December 4
Stop 9a and 9b. BOUTONNIÈRE DE OUAOUFENRHA Low road cuts along the road from Taroudant to Igherm northwest of Taguenza and close to the Boutonnière de Ouaoufenrha allow some insights into the lithologies of the Tabia Member of the Adoudou Formation. Parts of the Tabia Member close to the stop have presented poorly preserved ribbon-like carbonaceous structures that we interpret as possible vendotaenid-type fossil remains. Similar vendotaenids are primarily known from the terminal Ediacaran (Vendian) of the East European Platform, where they have been commonly regarded as index fossils of latest Proterozoic age. However, vendotaenids have also found in other regions where the
stratigraphic context is less striking. A restudy of the Proterozoic–Cambrian Nama Group has led to the discovery of vendotaenids in the Schwarzrand Subgroup, where they clearly appear within earliest Cambrian rocks (GEYER 2005b) according to the internationally accepted concept of the Precambrian–Cambrian boundary. Directly south of Ouaoufenrha (Wawfengha), the road cut on the western face exposes the Tabia Member of the Adoudou Formation in a carbonate-dominated facies with local developments of calcimicrobial mats and stromatolitic bioconstructions.
December 4
Stop 10. OUI-N-TATAYINE The huge ‘boutonnières’ of Ida ou Zedout/Ida ou Zekri (= ‘Igherm massif’) and Izaen are largely formed by rocks belonging to the upper part of the so-called Précambrien II-III (or ‘système de Tidiline-Lizate’; CHOUBERT 1963). This unit was originally interpreted by CHOUBERT (e.g., 1963) and others as composed mainly of flysch deposits which were formed between the orogenic events of the Anti-Atlasids (ca. 1,400 Ma b.p.) and the Marocanids (ca. 1,000 Ma). However, the older radiometric age determinations are now understood to be notoriously misleading. In addition, the many formally named rock units prove to be geographically restricted, and are difficult
to correlate between the basement massifs in the AntiAtlas. Assignment to the different stratigraphic levels such as the Précambrien I, II, II-III, and III and their subunits has depended on the degree of metamorphism rather than precise age determinations or lithology and sequence. THOMAS et al. (2004) suggested a new lithostratigraphic framework for the Neoproterozoic rocks of the Anti-Atlas orogen. Despite the relatively preliminary subdivision, the rocks assigned here to the traditional Précambrien II-III of CHOUBERT would belong to the ca. 650 to 800 Ma Anti-Atlas Supergroup in the concept of THOMAS et al. (2004).
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However, at least some parts of the Ida ou Zedout/Ida ou Zekri and Izaen massifs are problematical in their stratigraphic position. Particularly wide outcrops are formed by the Ida ou Gnidif Formation, which consists essentially of conglomerates. The boulders show a wide range of sedimentary and igneous rock lithologies, which prove a complex source area. East of Oui-n-Tatayine, the road exposes easily accessible, low outcrops of the Ida ou Gnidif conglomerates. Elsewhere, intercalations of green shales and sandstones and locally reddish shales are found. Tillites have been reported, as well, from the formation. These diamictites suggest that the rocks correlate with a widespread Neoproterozoic glacial event of ~700 Ma rather than the younger ~600 Ma old glacial period which is post-Pan-African. However, evidence for their
glacial origin requires further investigation. Halfway between Igherm and Oui-n-Tatayine, the road crosses variable levels in the Ida ou Kensous Formation. This stratigraphic unit consists mainly of reddish to purple sandstones with frequent layers of coarse-grained sandstones and conglomerates. Stratiform volcanogenic rocks such as rhyolitic ignimbrites and andesitic rocks, are intercalated. The Ida ou Kensous Formation is a disconformably bound unit which according to traditional views (e.g., CHOUBERT 1963) forms the oldest part of the so-called ‘Précambrien III’ in this region. This unit is, in fact, underlain by older rocks of the Précambrien III elsewhere. Its age is somewhat problematic because reliable radiometric datings of this unit are lacking.
December 4
ISSAFEN SYNCLINE The Issafen area in the center of the western Anti-Atlas is formed by a narrow but elongate, north–south oriented syncline between the large Precambrian Kerdous (in the west) and Izni (in the east) massifs. The syncline has a nearly vertical axis (Fig. 31). The core is formed by Cambrian deposits, and the youngest strata are Middle Cambrian. The Igoudine and Amouslek Formations are well exposed, and form nearly vertically dipping slopes on the eastern flanks of the syncline and slightly less steeply dipping strata on the western slope. A short stop will demonstrate the facies and transition from the Issafen into the Tazlaft and Tatelt Formations close to the type section of the Issafen Formation. Issafen Formation The Issafen Formation consists of grey, yellowish, reddish, and purple slaty shales and siltstones with only a few calcareous and/or quartz arenite beds in the lower part. Rubbly or nodular limestones (the so-called Cal-
caire scoriacé) primarily occur in the upper part. The lower part of the Issafen typically consists of slightly micaceous, finely laminated, light-colored shale and siltstone, often with a slaty cleavage. The fresh color of the rock is greenish to light yellow. It weathers to very light grey to white, sometimes with light yellow to light brown lines that mark calcareous laminae. The facies of the closely packed calcareous nodules with a shale ma-trix (i.e., „scoriaceous limestones“) is limited in the Issafen Formation to the western AntiAtlas. These calcaires scoriacé beds form the base and top and comprise thin intervals within the formation. Calcaire scoriacé has been interpreted as diagenetically altered beds of fine-grained terrigenous clastic and subordinate volcanoclastic sediment and calcareous interbeds (DE KONING 1957). However, a more probable origin lies in diagenetic redistribution of calcareous material within green, purple, and red siliciclastic muds and the growth of nodules within burrow-churned sediment (GEYER et al. 1995).
Fig. 31. West–east cross-section through the northern part of the Issafen; the more strongly upwarped eastern flank has nearly vertically dipping basal Cambrian units. Modified from CHOUBERT (1952).
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The type section (about 2 km south of Timghit) and other sections in the Issafen area show a monotonous, thick succession of green, slaty siltstone with occasional fossil remains (mostly trilobite sclerites), interrupted rhythmically by bioclastic limestone beds, which occasionally yield archaeocyathans. The Issafen area has the most fossiliferous successions of this formation known so far. Several sections were studied in the 1940s and 1950s, and a number of trilobite taxa have their type localities in the area, e.g. Bondonella typica, Neltneria jacqueti, Pareops transitans, Perrector pocteyae, Longianda termieri, Pseudosaukianda lata (Fig. 32), Gigantopygus bondoni (Fig. 32),
G. papillatus, and G. angustalatus. However, none of the sections have been measured precisely, and most of the samples were collected from float. In addition, due to the usually strong tectonic deformation fossil remains suffered generally a conspicuously distorted. Easily accessible outcrops of the Issafen Formation with relatively moderate tectonic deformation are located between Tamachourt and El Kesbet east of the road. The shales of the upper Issafen Formation is these sections are relatively fossiliferous with a moderately divers trilobite fauna. However, these sections were never measured accurately.
December 4
STOP 11. CENTRAL ISSAFEN SYNCLINE Two sections with overlapping stratigraphic intervals in the core of the Issafen syncline are located about 6 km south of Souk El Khemis (or El Kesbet, geological map sheet Tafraout 1 : 100,000). These sections show the uppermost Lower and lowermost Middle Cambrian as the youngest pre-Quarternary strata in the Issafen Syncline. The stratigraphically higher section is on the southern slope of the hill and within the ruins of Amadine (section IS I, Lambert coordinates 197/319.0). The stratigraphically lower section is along the southern slope of the hill 400 m to the north (section IS II, Lambert coordinates 197/319.4). Tazlaft Formation A Termierella latifrons faunal assemblage collected by HUPÉ (1953, section Timghit-1, sample F8, section Timghit2, samples F1, F2, F3, and F5) was reported to occur in the upper part of an extraordinarily thick Issafen For-
mation (370 m, HUPÉ 1959). This assemblage is referable to the Sectigena Zone in other sections. However, various authors (HUPÉ 1953, 1959; CHOUBERT 1952; DESTOMBES 1985; SIEGERT 1986) claimed that the „grès terminaux“ or „Asrir Formation“ [i.e., Tazlaft and Tatelt Formations in corrent terminology; LANDING et al. (2006)] was not present in the Issafen area. This created the myth of the „Island of Issafen,“ which was supposed to have existed during the latest Early Cambrian. However, lateral equivalents of the upper „Asrir Formation“ (i.e., the Tatelt Formation of LANDING et al. 2006) as well as the Jbel Wawrmast Formation were indeed deposited in the region as shown by GEYER (1990b) and HELDMAIER (1997). The basal 15 m of section IS II are part of the Aït Herbil Member of the Tazlaft Formation. A major stratigraphic gap that includes parts of the „Asrir Formation“ and the Jbel Wawrmast Formation is not present in the Issafen Syncline. These facts suggest a reevaluation of the strati-
Fig. 32. Examples for characteristic trilobites from the Issafen Formation in the Issafen Syncline. Left: Gigantopygus bondoni HUPÉ 1953; right: Pseudosaukianda lata HUPÉ 1953. Both latex casts of the holotypes. Museum Nationale d’Histoire Naturelle, Paris (MNHN R 50822 and MNHN R 80883). Scale bars equal 5 mm.
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thick interval can be identified as the Tatelt Formation. The division into a lower coarser unit and an upper finergrained interval suggests close similarities to the Akka Tzem Member and Aït Herbil Member of the Tazlaft Formation. Thus, the samples with Termierella latifrons now most probably came from the lower Tazlaft Formation. The uppermost 16.5 m of the Tazlaft Formation consist of predominantly consist of gray to dark green, mediumand coarse-grained arenites. Their composition ranges from well-sorted quartz arenites, lithic arenites, and feldspathic arenites to arkoses. The feldspar content is generally high (10–30 %); euhedral feldspars are common. Even twinned feldspars are present. The 1 to 2 m-thick beds are often graded and occasionally have a 10 to 15 cm-thick conglomerate at their bases (clasts up 5 cm in length). At 12.55 m a single green shale layer was observed (< 1 cm). Due to intensive weathering, sedimentary structures have not been observed. The thick bedding, basal conglomerates, and finingup trends within 1 to 2 m-thick beds demonstrate high energy deposition in a unidirectional flow regime, most likely in a fluvial or deltaic system or in a tidal environment. Occasionally well-rounded quartz grains indicate an intensive reworking as an indicator of fluvial transport or reworking at the shoal. The presence of angular quartz and euhedral feldspar as well as the variable composition between the beds indicate a heterogeneous source area. Tatelt Formation
Fig. 33. Detailed Tazlaft through Jbel Wawrmast Formations of sections IS I and IS II in the central Issafen Syncline, illustrating grain-size profile, color changes, and lithostratigraphic succession. Thickness (left margin) in meters from the base of the measured section. See Fig. 27 for legend. Slightly modified from HELDMAIER (1997).
graphic nomenclature applied at other measured sections in the region (HUPÉ 1959; SIEGERT 1986). GEYER’s (1990c) reference to two measured sections in the Issafen with a total formation thickness of 370 m, which was based on data from HUPÉ (1959). In HUPÉ’s (1959) report, the Issafen Formation is overlain by the Asrir Formation. The fossiliferous samples with Termierella came from the upper 100 m of the „Issafen Formation“. SIEGERT (1986, fig. 92) showed a 340 m section for the Issafen Formation. Above the highest red shale interval he reported a 111 m-thick unit composed of finegrained, siliceous sandstones (the lower 79 m) and a shale interval (the upper 32 m). The next overlying, 46 m-
The Tatelt Formation in the Issafen syncline is dominated by yellowish-green, fine-grained sandstones. They are interrupted by single, usually coarser-grained marker horizons. The basal unconformity of the 43.3 m-thick Tatelt Formation (and base of member 1 of the formation) in the center of the Issafen syncline is overlain by an up to 25 cm-thick conglomerate composed of vein quartz pebbles that float in a fine-grained matrix. The bed is cemented with carbonate at the top. The lowest yellowish-green, fine-grained sandstones in the section appear above the conglomerate, but dominate the Tatelt Formation. Within member 1, two bluish horizons, locally with calcareous nodules, are the lateral correlatives to the shoaling-up sequences ob-served in the eastern Anti-Atlas and other areas (LANDING et al. 2006). An almost 40 cm-thick bluishgreen interval (17.8 m) represents the lower horizon of the two carbonate horizons of member 1 [see LANDING et al. (2006) for informal members 1–3 of the Tatelt Formation]. The base of member 2 is a 5–10 cm-thick, coarsegrained, arkosic tuff that is composed predominantly of euhedral feldspars. Similar to other sections, member 2 consists of two shoaling-up sequences. The lower parts
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Fig. 34. Youngest Cambrian strata in the Issafen Syncline: Tatelt and base of Jbel Wawrmast Formation at the hill with the ruins of Amadine, ca. 6 km south of Souk El Khemis (El Kesbet). View towards south. Unterlying Tazlaft and Issafen Formations to the right. In the far distance the Lie de vin and Adoudou Formations and basement rocks on the western flank of the syncline.
are fine-grained and without recognizable sedimentary structures. Lamination increases upward and the rocks become bluish in color and slightly more calcareous, then grain-size increases to coarse-grained, and evidence for such higher water energy conditions as megaripples, occur. Both, the nodular carbonate and the arkose, are interbedded with a bluish-green interval (22.9–23.7) with abundant fine-grained quartz. These horizons also crop out at the Souk El Khemis (El Kesbet), 6 km to the north, where samples (IS-S1 and IS-2 of HUPÉ 1959, and GEYER & HELDMAIER, unpubl.) contain a distinctive Hupeolenus Zone fauna. Between 23.7 and 32.65 m, the fine-grained yellowish-green sandstones are structureless. Only the top meter is laminated; it consists of 0.5 to 1 cm-thick beds, and occasionally shows micro-crossbedding. A yellowish white, medium- to coarse-grained arkose overlies the laminated sandstones (32.65–36.6 m). The arkose thickens from 1.8 m (section IS II) to 3 m (section IS I) within a distance of only 400 m. The arkose (IS I-13.5 m) is overlain by yellowish-green, fine-grained sandstones which occur in 3–15 cm-thick beds, and include bluish to gray-green patches. A calcareous horizon is intercalated at 21.6 m. Between 23.25 and 24.9 m, megaripples dominate in the lower half and dunes in the upper half. The lower half is a gray, calcareous, fine- to medium-grained arenite, the upper half a gray, fine-grained sandstone to siltstone with some calcareous patches. Member 3 consists predominantly of fine-grained sandstones. An intercalated, laminated and microcrossbedded interval indicates a zone of possibly shallower water. Above it is a shoaling-up sequence similar to that in member 2. The background sedimentation of finegrained, yellowish-green sandstones indicates a relatively low energy depositional setting. Member 2 and member 3 contain four similar cycles, which suggest a shoaling-
up from a low- into a high-energy environment. Above the base of member 3 (24.9 m) yellowish-green, finegrained sandstones with a grayish cast dominate the section. Between 28.4 m and 30.8 m, lamination is more frequent, and small-scale ripples and microcross-bedding are observed locally. The laminated interval ends with a 50 cm-thick, more massive bed. Above it, the rocks are characterized by dark, unidentified minerals. They are massive (34.25–34.8 m). An interbedding of yellowishgreen, fine- and medium-grained sandstones terminates the formation. Trough crossbedded beds reach a thickness of 40 cm. Jbel Wawrmast Formation Fine-grained, locally calcareous, bluish-green sandstones deposited on an erosional unconformity on the Tatelt Formation form the base of the heterolithic Jbel Wawrmast Formation. The erosive basal unconformity of the Jbel Wawrmast Formation is overlain by a heterolithic interval (36.6–39 m) with quartz arenites of different grain sizes and a 60 cm-thick arkose. The variable composition and ripple marks indicate higher energy condition during deposition by comparison with the overlying greenish, fine-grained sandstones. The overlying monotonous yellowish-green, fine-grained sandstones have a coarse-grained base, which fines up into fine-grained sandstone (39–59.2 m). Up to 49.3 m, the interval is laminated and comprises 4–15 cm-thick beds. Limestone nodules are intercalated at 41.6 and 41.9 m. At 48.35 m, the base of a color cycle is bluish-green (48.35–49.3 m) and its lower part has an increased lamination and contains calcareous patches. A 40 cm purple interval follows and is overlain by a 45 cm-thick, calcareous, fine- to coarsegrained sandstone with megaripples. An additional well preserved shoaling-up sequence
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(50.3–61.3 m) begins with a basal, non-laminated interval followed by a laminated and massive bedded interval, and is terminated by a coarser-grained, large-scale, trough cross-bedded interval, which indicates the shallowest depositional environment. The top of the massively bedded interval (59.2–61.3 m) is slightly coarsergrained with large scale trough cross-bedding. An overlying thin glauconite sand is followed by a pink siltstone (61.3–61.4 m), and reflects an extraordinarily reduced sediment accumulation rate. Condensation is indicated by a 2 cm-thick glauconite sand followed by a pink siltstone at its top. Between 61.4 and 70 m, yellowishgreen, fine-grained sandstones contain occasional smallscale, trough cross-bedding and calcareous patches. A
brachiopod hash layer is inter-calated at 62.6 m. Above 70 m, the sedimentary rocks comprise 15 to 20 cm-thick beds and are purple up to the top of the section. Phosphatic nodules occur between 70 and 70.4 m. At 71.3, 71.4, and 71.5 m, single shell hash beds are intercalated and form part of a 1.2 m-thick massive shell hash bed with intercalated 5 to 15 cm-thick intercalated siltstone and sandstone intervals at 72 m. Some of the bases of the shell hash beds locally show „giant traces.“ The uppermost limestone-dominated interval is divided by a 10 cm-thick purple shale and contains stromatolite mats every 5 to 10 cm. Within these limestones a trilobite fauna, dominated by Kingaspidoides borjensis probably belongs to the Cephalopyge notabilis Zone.
December 4
STOP 12. SOUTHERN ISSAFEN SYNCLINE Well-developed archaecyathan bioherms appear in the upper part of the Cambrian succession in the southern Issafen Syncline. A particularly spectacular example is visible on the western side of the road south of Welldeveloped archaecyathan bioherms appear in the upper part of the Cambrian succession in the southern Issafen Syncline. A particularly spectacular example is visible
on the western side of the road south of Zaouïe n’Ait Haroum, close to Asgoun Doû Tourirt, where a large, wedge-shaped build-up gives insights into the complex structure of those Early Cambrian bioherms. It is partly composed of archaeocyathan buildups and isolated cups, with an even higher percentage of the carbonate buildup constructed by calcimicrobes which fill part of
Fig. 35. Section close to Asgoun Doû Tourirt, southern Issafen Syncline. View towards south shows Issafen Formation (dipping toward west) with typical shales and a well-developed archaecyathan bioherm that forms the top of the hill. The Lower boundary of the bioherm is trough-shaped, but the archaeocyath-microbial bioconstructios pass laterally into limonitic peloidal and scoriaceous limestones. Blocks on the slope below the bioherm in the photo are fallen from the reef limestone complex.
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the space between the archaeocyathans. The spatial composition here indicates that the actively growing part of the bioconstruction was low, probably only a few centimeters high, and probably forming a gentle barrier in the shallow, moderately wave-agitated sea. Niches between archaeocyathans and calcimicrobial carbonate ledges were filled with skeletal remains of such mobile and semisessile benthic animals as hyoliths, trilobites, and brachiopods. The Issafen archaeocyathans include nine species (partly reported with in open nomenclature), which appears to indicate a correlation into the lower Botoman
Stage of the Siberian Platform. (The Siberian archaeocyathan succession serves as a quasi-standard for the global archaeocyath biostratigraphy.) The Issafen archaeocythans (DEBRENNE & DEBRENNE 1995) include: Coscinocyathus amagurensis (DEBRENNE 1964), Afiacyathus undatus DEBRENNE 1964, Archaeopharetra sp., Chouberticyathus clatratus DEBRENNE 1964, Erismacoscinus calathus (BORNEMANN 1887), Bottonaecyathus sp., Dictyocyathus verticillus (BORNEMANN 1887), Dokidocyathus rarus DEBRENNE 1964, and Protopharetra sp. aff. P. polymorpha BORNEMANN 1887.
December 4
STOP 13. WEST OF IMITEK The paved road from Tata to the Issafen syncline crosses west of Imitek the upper Adoudou and Lie de vin Formations in the region west of Imitek. The facies of these formations in this region is dominated by an alternation of calcimicrobial dolostones and limestones and purple shales. The Hercynian orogeny has created conspicuous slump folds which will be studied from the road.
Lie de vin Formation The rock unit today designated the Lie de vin Formation was recognized in the region of Tata and, particularly in this area, as the typical „purple slates“ (CHOUBERT 1942). Thus, this stop west of Imitek gives a perfect impression of the original concept of the formation. The lateral facies change in the western Anti-Atlas is considersable and illustrated in Fig. 36. On this southern slope of the western Anti-Atlas, the sections are dominated by purple and reddish shales (argillaceous pelites). The thickness is about 650 m. The term lie-de-vin refers to the typical purple color of the shales that resembles the color of a French Burgundy wine. The upper third of the formation features a striking increase of dolomites. A massive carbonate unit termed the ‘Barre de Tata’ (Tata ledge), which is difficult to identify in the north of the Anti-Atlas, is particularly evident. In addition, earlier
workers distinguished a similar carbonate unit as the ‘Akka Irhen bed’ at the top of the formation. The carbonate beds are important for understanding the formation’s deposition and stratigraphy. In fact, the Barre de Tata was demonstrated to step up in the section, and it indicates that the Lie de vin deposition was diachronous. The increase in siliciclastic content above that of the Adoudou Formation reflects a relative regression visible in most section. However, this is interpreted as a result of a seaward tilting of the Anti-Atlas margin at the end of the Adoudou deposition. This epeirogenic event initiated a marked change in deposition: the tilting, or flexure is indicated by a significant regression and production of an erosional contact with the underlying units in eastern areas of the Lie de vin deposition. However, distally the formation features a change from peritidal carbonate into sublittoral shale–biohermal limestone sedimentation with an increase of terrigenous detritus. The region south of the Issafen Syncline and west of Imitek is an example of fine-grained, siliciclastic-dominated input, whereas the earlier visited Tiout section exemplifies the alternating carbonate–shale lithofacies which appears to reflect some combination of changes in siliciclastic input, salinity, or climate or small-scale sea level changes.
December 4 STOP 14. HASSI BRAHIM The road from Imitek to Tata passes a section ca. 5 km west of Hassi Brahim, which exposes a complete succession from the top Igoudine Formation into the Jbel Afraou Formation. This Hassi Brahim–West section (GEYER & LANDING, unpubl.) shows a marginal develoment of the Lower Cambrian strata with the Amouslek Formation largely developed as dolostone–shale cycles and its lower and upper boundaries difficult to identify. The
Issafen Formation is suprisingly thin compared with neighbouring sections the southern Issafen Syncline. The main Hassi Brahim section is located north of the road from Tata to the Issafen area, and 5 km to the WSW of Tata (Lambert coordinates 240/304, geological map sheet Akka–Tafagount–Tata 1 : 200,000). This section shows a complete succession from the Igoudine Formation (Lower Cambrian) through the Middle Ordovician.
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It shows a development from a marginal facies of the Igoudine, Amouslek, and Issafen Formation to a progressively relatively distal facies if the Jbel Wawrmast and Jbel Afraou Formations. All of these units are conspicuously thinner by comparison with sections in the axial part of the Anti-Atlas. The upper part of the section has been described in detail by HELDMAIER (1997). This section stretches along a N–S oriented creek that crosses
the plain at the southern margin of Jbel Azrouisse. The base of the section is located in a gorge at Jbel Azrouisse. In the lower part of the section, the Igoudine and Amouslek Formations are developed as alternations of carbonates and shales. Massive, shallow-water carbonates, often as dolostones, dominate the succession. Fossils are absent. The Issafen Formation consists of slaty grayish shale with only sparse trace fossils.
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Fig. 36. Facies and lateral development of the Lie de vin anf lower Igoudine Formations central part and southern flank of the western Anti-Atlas. Note easily recognizable increase in calcareous rocks upsection in the western sections and the persistence of a maximum of calcareous deposition in middle member of the Lie de vin Formation (identified as the Barre de Tata on the southern flank of the Anti-Atlas). After unpublished data of FORTEMS (1963, 1964); slightly modified from DESTOMBES (1985, fig. 11).
Tazlaft Formation Hassi Brahim is the type locality for the western facies of the Tazlaft Formation. Within an interval of about 20 m, isolated quartz arenite beds as well as series of amalgamated quartz arenite beds are intercalated with red shales. Most of the sandstones are structureless, but some show HCS. The shales between the beds are locally burrow-churned.
The base of the Tazlaft Formation is marked by a change in color from red to green and a change in lithology from shales to medium-grained lithic and quartz arenites. The Akka Tzem Member (77 m) can be divided into two parts: The lower part (64–97 m) is dominated by greenish to yellowish-green, medium-grained arenites. A roughly one meter-thick horizon at 71.5 m is intensely red stained. The bedding thickness varies from several centimeters to 1.5 m. Shales are occasionally preserved as laminae on bedding surfaces. Shales form two intervals up to 1.5 m (89.5–91 m) and 2 m (94–96 m) in thickness about in the middle of the member. The upper shale interval marks a slight change in lithologic composition of the arenites (i.e., the upper half of the Akka Tzem
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Fig. 37. Tatelt–Jbel Wawrmast Formation contact in the Hassi Brahim section. Photo shows steeply dipping quartz arenites of the top of the Tazlaft Formation with sedimentary structures such as wave and interference rippes overlain by condensed calcareous bed with accumulations of brachiopod valves (Trematobolus superbus) and few trilobites of the Jbel Wawrmast Formation at right lower corner.
Member, 96–142.5 m). Feldspar becomes abundant and the color changes to yellow with this facies change. A 2 m-thick, bioturbated, calcareous shale interval contains limestone nodules at 128 m. Conglomeratic horizons are intercalated in whitecolored, well washed quartz arenites and conglomeratic quartzites of the overlying member (11.5 m). Their poorly sorted components consist almost entirely of vein quartz pebbles. They are generally 0.5 to 1 cm in diameter, but reach up to 7 cm in length. Low-angle cross-bedding is often observed. The conglomerates in the lower half form several centimeter- to decimeter-thick beds, but amalgamate to form a 4 m-thick conglomerate at the top. The fine-grained beds between the conglomerate horizons are locally bioturbated. The Aït Herbil Member (154–251m) is a relatively monotonous, yellowish sequence of interbedded, finegrained siliceous sandstones and feldspathic arenites (20-50 cm). The interval from 155 to 190 m is intensively bioturbated by organisms that created horizontal and vertical traces. A 3 m-thick, white, well washed quartz arenite is intercalated at 176 m. Skolithos is abundant in the top 50 cm of these quartz arenites. Between 190 and 193 m the arenite beds become more abundant and thinner bedded. The overlying interval is a 10 m-thick, green colored, massive, homogenous quartz arenite. It is capped by 1.5 m of fine-grained, red siliceous sandstones. The uppermost strata of the formation are again yellowish, fine-grained, siliceous sandstones interbedded with medium-grained arenites. The arenites only occur in the 221–235 m interval. The intercalation of siliceous arenites in the upper part of the shale-dominated Issafen Formation indicates an increase of coarser-grained terrigenous input onto the shelf. This is probably the result of a prograding
facies shift to the west of the Akka Tzem Member. The appearance of HCS suggests further additional shallowing of the depositional setting into the lower shoreface. In the Akka Tzem Member, only the thin shale interval between 127.5 and 129 m can be interpreted as deposited in a marine environment. This is based on the bioturbation and the greenish color of the rocks. The latter indicates reducing conditions, at least after shallow burial. Compared with the Akka Tzem section 40 km to the east, the arenites are generally finer grained. They are relatively well-sorted, and apparent sedimentary structures do not indicate strong current or wave activity. This indicates a redeposition of well-sorted sands into a deeper part of the basin. A source area for the material was possibly located to the east where well-sorted fluvial (intertidal?) sands were transported directly to the west. The slight increase in feldspar content at about the middle of the Akka Tzem Member could be the result of erosion of basement or earlier deposited feldspathic sandstones. The intercalated bioturbated, calcareous shales represent a period of lowered siliciclastic input. A general shoaling-up sequence similar to other western sections can not clearly be identified. The middle member reflects the shallowest environment of the Tazlaft Formation. The sandstones were well sorted by wave activity, and were bioturbated in a marine setting. The gravel is extraordinarily well rounded and spherical. Much larger clasts (>4 cm in length) are elongated blades. The shape of the components as well as the low-angle foresets in the conglomerate (which dip towards the west) indicate beach deposits that possibly accumulated in a supratidal setting (beach berm). The Aït Herbil Member of the Tazlaft Formation was deposited in a marine environment, as indicated by abundant bioturbation. A distal deltaic or tidal-influenced
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setting are both appropriate. The three massive sandstone complexes that interrupt the rhythmic succession are noteworthy. The well-washed quartz arenite with Skolithos at its top is typical of shallowing-up shoreface deposits. The second sandstone complex contains a color cycle: a green arenite with a red (oxydized) shale top. This again reflects a shoaling upsection. Tatelt Formation At 252 m, at section HBR II (1.5 m below the base of section HBR), the 26 m-thick Tatelt Formation conformably overlies the Tazlaft Formation. The change is marked by the loss of the siliceous sandstone interbeds at the top of the Tazlaft Formation and the change in color from light yellowish-green to a medium-green. A monotonous, indistinctly laminated, fine-grained sandstone represents the base of the Tatelt Formation. A single horizon at 1.9 m contains up to 20 cm-long and 3 cm-thick carbonate nodules. Between 5.7 and 6.9 m this interval is slightly more arenaceous and coarser-grained. Brachiopods occur on a bedding surface at 6 m. A 5 cmthick, medium-grained, slightly calcareous volcanoclastic arkose occurs in a 16 cm-thick, yellowish-green shale interval. Above the shale, a yellowish gray to black, fineto medium-grained arenite (7.1–7.7 m) contains subvertical and horizontal burrows at the base. Between 7.7 and 9.15 m, two fine-grained sandstone beds of a meter in thickness show a progressive upward fining into shales, which are overlain by 1 cm to 5 cm-thick beds of strongly laminated, fine-grained sandstones with occasional micro-crosslamination. Both intervals are mediumgrained. A 20 cm thick, medium-grained arkose with orange to yellow color terminates this interval. Between 11.3 and 12.5 m, several 30 cm-thick, mediumgrained arenite beds fine upward into shales. The same trace fossils observed at 6 m are present at the base of the highest and thickest fining-up bed. A homogenous, indistinctly laminated, fine-grained, medium-green sandstone overlies the medium-grained unit (12.5– 15.8 m). Between 15.8 and 16.75 m, these sandstones are strongly laminated. Two 20 cm-thick, fine-grained, white quartz arenites are intercalated at the top of the laminated units, and the latter are overlain by an indistinctly laminated, yellowish-green, fine-grained sandstone. Between 17.35 m and the top of the formation (24.65 m), fine- to medium-grained, white quartz arenites dominate the formation. Fining-up beds are observed between 19.5 and 20.25 m,. Fine-grained, medium-green arenites are intercalated between 21.7 to 22.15 m and 22.45 to 22.65 m. They include thin medium-grained quartz arenite layers which contain wave ripples at their top. The uppermost interval (22.65–24.65 m), which is composed of 20 to 30 cm beds, contains 1 to 8 cm-thick shale layers at the
Fig. 38. Tatelt and lower Jbel Wawrmast Formations of the Hassi Brahim section, illustrating grain-size profile, color changes, and lithostratigraphic succession. Thickness (left margin) is in meters from the base of the Tatelt Formation and does not refer to the base of the measured section. See Fig. 27 for legend. Slightly modified from HELDMAIER (unpublished).
top of the beds. In addition, it contains medium trough cross-bedding. The top of the formation is an erosional surface at the base of the Jbel Wawrmast Formation.
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In terms of depositional environment, the Tatelt Formation shows a general tendency of low-energy deposition of homogenous, fine siliciclastics in member 1, a larger variety of different sediments in member 2, and predominantly coarse-grained siliciclastics in member 3. Most characteristic for the formation at Hassi Brahim are up to 1 m-thick, fining-up beds in members 2 and 3, which in most cases suggest a deposition by turbidity currents. They indicate a deposition on or distal to a slope. As in other sections (Aguerd, Oued Boutergui, Amouslek, Akka Tzem) a water-lain tuff marks the base of member 2. No evidence for wave deposition is determinable in member 1, which suggests a deposition below the stormwave base. Current activity increases in member 2, as indicated by the occurrence of turbidites and more intensive lamination. In member 3, the sediments are better sorted and reworked. The average grain-size increases and turbidites still occur. Toward the top of the formation, wave ripples, trough cross-bedding, and reworked shale clasts are preserved. These observations point to a shallower setting than for the member 1 and 2. In summary, the Tatelt Formation at Hassi Brahim gradually shoals upward. While member 1 seems to have been deposited under relatively quiet marine conditions below the storm-wave base, members 2 and 3 show an increased input of coarser siliciclastics that were predominantly deposited as turbidites. The sediments of member 3 are coarser-grained, enriched in quartz, and show wave ripples and reworked shale clasts at the top, which suggest the shallowest environment of the Tatelt Formation in the Hassi Brahim section. The turbidites indicate either a position on the slope or at the slope base. However, the formation shows almost all of the characteristic marker beds which are observed in other sections, and which are interpreted as the result of minor sea-level fluctuations. Jbel Wawrmast Formation The Jbel Wawrmast Formation (92.85 m-thick) overlies the quartz-arenitic top of the Tatelt Formation unconformably with a basal, fossiliferous, calcareous conglomerate (24.65 m), which is predominantly composed of several centimeter-long, laminated shale clasts. Within 40 cm the conglomerate fines into fine-grained sandstones which characterize the entire formation. The basal 7.45 m (up to 32.1 m) of grayish green sandstones are intensively laminated, and intercalated with isolated, 3– 5 cm, slightly coarser-grained sandstone beds. Between 32.1 and 53.3 m, the sandstones are only slightly laminated, and medium-green in color. Slightly coarsergrained, 3–5 cm-thick sandstones, as well as two single carbonate nodules (40.1, 45.35 m), and two, ca. 10 cmthick conglomeratic horizons are intercalated. The lower
of these conglomeratic beds (34.10 m) consists of flat shale pebbles embedded in a calcareous matrix. In the upper bed (42.45 m) the shale clasts are much more rounded and contain additional quartz clasts. The components of this bed float in the matrix. At 53.2 m, the color turns abruptly to purple. This 10 cm-thick, purple interval is overlain by 70 cm of a yellowish arkose with a basal lime mudstone layer. The carbonate content decreases towards the top of the arkose, and this bed becomes more quartzose. Between 54 and 56.8 m, the rock is bluish-green. The basal meter of the interval, is predominantly laminated. A calcareous nodule horizon is intercalated about at the middle of this bluish-green interval. Between 56.8 and 61.05 m the rocks are bright purple. Black, elongated calcareous mud-stone nodules occur at 57, 57.6, 57.7, and 57.8 m. From 59.05 to 61.06 m, fossiliferous, calcareous, fine-grained sandstones and overlying pack- and wacke- to grainstones up to 50 cm-thick alternate with thin shale beds. The shell hash has an edgewise orientation particularly in the thicker limestone beds. Obolellid brachiopods (Trematobolus superbus) dominate the basal fine-grained sandstones, whereas trilobites dominate the faunas in the pure limestones. Due to the strong hematite cement, fossil recovery is poor. The massive limestone at the top is overlain by 45 cm of grayish-green, fine-grained sandstones, which have limestone nodules at the top. This interval (53.2-61.5 m) represents the top of the Brèche à Micmacca Member. Overlying, yellowish-green, fine-grained sandstones characterize the deposits up to the top of the formation. The lower 14 m (up to 75.8 m) are completely structureless. At 75.8 m, fossiliferous grainstone nodules are intercalated, overlain by a 10 cm-thick grainstone bed at 76.8 m. Above this limestone, 5 to 15 cm-thick, intensively laminated intervals are rhythmically intercalated in 2 to 30 cm alterations. Above 85.25 m, the rocks are laminated, but include isolated, 2 to 10 cm-thick, non-laminated intervals. The lamination ends at 91.5 m. Between 101.5 and 106.55 m, occasional calcareous nodule layers are associated in a laminated sandstone facies. At the top of this interval is a lenticular limestone bed (wacke- to grainstone) of up to 60 cm in thickness. This lens pinches out over a distance of about 50 m. The ferruginous shell hash contains shale clasts up to 5 cm in length and shows megaripples. Fossils remains consist predominantly of phosphatized trilobite sclerites, most probably indicating the Ornamentaspis frequens Zone. This interval represents the top of the Tarhoucht Member (introduced by GEYER & LANDING, this volume). Units above the Jbel Wawrmast Formation The rocks above the Jbel Wawrmast Formation have similar grain-sizes, but prominent lamination and thicker
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bedding reflects deposition under higher energy conditions. Between 107.15 and 117.75 m, well laminated beds are intercalated. Higher strata show a cm-scale bedding, are locally laminated, and are overlain by an intensively laminated interval of 10 to 20 cm-thick sandstone. Between 126.55 and 149.7 m, the bedding thicknesses progressively increase from 1 to 10 cm (up to 130 m), to 5 to 15 cm (up to 135 m), to 10 to 30 (up to 144 m), and 10 to 50 cm (up to 149.7 m). The 10 to 15 cm-thick beds, particularly show trough cross-bedding. Above 149.7 m, the thick beds disappear, and 0.5 to 5 cm-thick beds are intercalated in intervals of 30 to 70 cm. Thin, 0.5 to 1 cm-thick sandstone intercalations occur above 156 m every several meters. Calcareous, fine-grained, 1–4 cm-thick are sandstones are intercalated between 161 m and about 216 m. These beds often form distinct pairs or triplets, while the stratigraphic distance bracketed by these grouped intercalations ranges from 1 to 3 m, and some-times up to 10 m. Calcareous nodules are rare. Kymataspis sp. aff. K. arenosa, Ornamentaspis sp. indet., and ‘Parasolenopleura’ sp. indet. characterize a Kymataspis arenosa Zone assemblage in this part of the section. At the top of the section (above 216 m), bundles of four, six, or sometimes even more than then calcareous
sandstone beds occur at intervals of 5 to 40 cm. Calcareous nodules are often associated. These bundles are 2 to 5 meters apart, but the distances separating them decrease up to 276 m. Sandstone intercalations are rare between 276 m and the top of the section (325 m). The fauna consists of the several trilobites (Fig. 39) such as Badulesia tenera, Kymataspis sp. aff. K. arenosa, and Paradoxides (Eccaparadoxides) sp., the brachiopod Brahimorthis antiqua, and echinoderm debris, and indicates deposition in the Badulesia tenera Zone. In summary, the basal unconformity of the Jbel Wawrmast Formation at Hassi Brahim is overlain by prominently laminated sandstones, which are themselves overlain by non-laminated deposits that indicate decreasing energy in the environment. The red-colored top of the Brèche à Micmacca Member has a thin purple horizon overlain by an arkose. This arkose is followed by a bluish-green interval with rare sandstone intercalations before carbonates appear in purple shales. The bluish-green color confirms a second shoaling-up sequence within the Brèche à Micmacca Member, which is overprinted by another cycle with a short period that generated the carbonate beds. The Brèche à Micmacca Member is followed by a lower-energy facies characterized by intercalated sandstones. The occurrence of carbonate nodules
Fig. 39. Characteristic trilobites of the Badulesia tenera assemblage in the Hassi Brahim section. Left: Badulesia tenera (HARTT 1868), complete carapace of young individual with incompletely developed ridges on the fixigenae. Right: Ctenocephalus antiquus THORAL 1946, more-or-less complete carapace. Note preserved canals in the tubercles. Both specimens from the assemlage at m 306 in the Hassi Brahim section. Coll. Institut für Paläontologie, Universität Würzburg.
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introduces the Tarhoucht Member (see GEYER & LANDING, Ediacaran-Cambrian depositional environments and stratigraphy, this volume), which is bounded by a thick (70 cm), purple, megarippled shell hash bed at the top. The sediments above the Jbel Wawrmast Formation do not show a change in grain-size, but lamination and particular bedding thickness reflect a progressive facies change during a period of higher energy conditions. Although a direct lithologic correlation to the Goulimine Formation or Jbel Afraou Formation is impossible the in-
terval was apparently deposited in the same shallow marine period. This shallower interval is followed by structureless, yellowish, to bluish-yellowish-green, fine-grained sandstones which represent distal shelf deposits. Fossils ca. 40 m (201 m) above the top of this period indicate the Kymataspis arenosa Zone. Above 216 m, these calcareous sandstones become thicker, calcareous nodules occur, and the fossils belong to the Badulesia tenera Zone.
December 5 STOP 15. ICHT A stop at a low road cut east of Icht (close to Foum el Hassane) demonstrates the Lower Devonian in the western Anti-Atlas. The exposures include a variable succession of limestones, shales and sandstones. The limestone beds are bioclastic and locally yield a diverse fauna dominated by trilobite sclerites. Trilobites attain a spectacularly large size, and are examples for the famous, often bizarre Devonian trilobites of Morocco, which generally come from the eastern Anti-Atlas. The section exposes the Assa Formation (Lochkovian?–middle Pragian) (largely forming the slope north of the road) and the overlying Merzâ Akhsai Formation (Pragian–lowermost Emsian; = Middle to middle to upper Siegenium in Rhenish stratigraphy). The overlying Ouin-Mesdour Formation (earlier regarded as part of the „Assise d’el-Ansar“, HOLLARD 1963, 1967; lower Emsian– Zlichovian) forms the hamada-type plain south of the road. The Assa Formation (from „Assise d’Assa,“ HOLLARD 1963) consists in the Foum el Hassane area and westward of an upward succession: 1) basal limestone beds, 2) middle shaly member, and 3) upper sandstone beds which are traditionally termed Rich 1 (used for the entire lithostratigraphic unit by HOLLARD 1963). Eastward, this tripartite division is difficult to recognize, and the formation was thus designated the Oued-el-Mdâouer Formation. The three members of the Assa Formation record a regressive development in depositional environment although the water depths apparently did not change dramatically. The basal limestone beds (20–100 m thick) show a change from crinoidal limestones with reddish intercalations to marly limestones and local nodular limestones, that passing gradually upward into claystones. The middle member consists mainly of sandy siltstones with layers of phosphatic nodules. Sandstone beds increase in number and thickness up-section, and finally develop into the Rich 1 sandstones, which are generally massive sandstone beds with scattered brachiopod valves and even brachiopod coquinas. Intercalations of oolitic layers and phosphatic nodules occur
at the top. The thickness of the Rich 1 sandstone beds varies between 20 and 50 m. This facies and faunal development records a change from a Hercynian to a Rhenish lithofacies, and from a neritic into a shallow marine biofacies. The brachiopods include the Pragian index fossil Platyorthis hollardi and Dixionella assaensis (JANSEN 2001, BECKER et al. 2004). The Merzâ Akhsai Formation (also spelled MerzâAkhsaï; from „Assise de Merzakhsai,“ HOLLARD 1963, 1967; concept modified by HOLLARD 1981) has its name from the ridge termed Jbel Merzâ-Akhsai located ca. 20 km to the south of the Stop 15 locality (Fig. 40). The formation has the same type of shallowing upward development. Basal limestone beds, a middle shale member, and upper sandstone beds (termed the Rich 2) are distinguished. The lower, 10 to 20 m-thick „Hercynian“ carbonates are bluish-grey crinoidal limestones of a variable microfacies with scattered reef-building organisms and pelagic faunal elements. Also present are numerous orthocones, which had a benthic mode of life and thus can not be attributed as pelagic fossils. Some beds include large sclerites of trilobites with brachiopods and conodont elements (upper Pragian with the co-occurrence of Latericriodus steinachensis and Caudicriodus curvicauda; BECKER et al. 2004). These basal limestone were originally assigned to the Assa Formation but later assigned to the Merzâ Akhsai Formation by HOLLARD (1981). The middle member consists of shales, siltstones and some sandstone beds with a thickness of up to 100 m. Phosphatic nodules are rare. Depositional environments obviously included poorly oxygenated oozes on the seabottom, and prevented the development of well-developed benthic communities. The Rich 2 sandstones are mostly thick-bedded, and include siltstone interbeds. Some brachiopod coquinas can be found in the upper part. The thickness is up to ca. 50 m. The age of earliest Emsian (= middle to late Siegenium in Rhenish stratigraphy) is indicated by a brachiopod assemblage with Filispirifer merzakhsaiensis („Acrospirifer fallax“), Ardu-
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Fig. 40. Lower Devonian outcrops along the main road from Foumel-Hassane to Tata east of Icht. Photo shows sandstones in the upper part of the Merzâ Akhsai Formation and the overlying Rich 3 in the distance. The mountain range in the background is Jbel Merzâ-Akhsai.
spirifer maroccanicus (JANSEN 2001). Other spectacular brachiopods include Mclearnites cf. cherguiensis and Iridistrophia sp. The lower surfaces of the beds south the road show an impressive series of rock engravings. Numerous animals such as rhinos, capricorns, jackals, and probably
antelopes, most of which are now unknown from the region, are depicted. However, symbolistic signs are present as well. Two types of engaving can easily be distinguished. They represent distictly different periods of its production.
December 5 STOP 16. ROADCUT EAST OF TIMOULAYE IZDER The main road from Foum-el-Hassane to Bou Izakarn passes a gap in the cuesta ca. 2.5 km to the east of Timoulaye Izder. This ridge is, in fact, the ‘First Bani’ (Fig. 41). Bani is the local term for the extended and relatively narrow ridges that envelop the southern Anti-Atlas and are separated by wide plains. Tabanite Group This First Bani is formed by the Middle to Upper Cambrian Tabanite Group (here with the Rich Khlifa, Bailiella, and Azlag Formations) and the basal Ordovician. The western slope exposes green, bioturbated, argillaceous sandstones of the Bailiella Formation. The best-exposed unit is the massive and thick-bedded quartz arenites of the Azlag Formation, which forms the crest of the ridge and is particularly rich in Skolithos tubes at this locality.
Goulimine Formation A brief stop ca. 1 km to the west will show an outcrop of the rarely exposed Middle Cambrian Goulimine Formation close to its type locality. Cliff-forming, coarse-grained quartz arenites and cross-bedded sandstone are best exposed. The sedimentary inventar includes low-angle trough cross-sets, current ripples, and small shale clasts. Poorly preserved Skolithos tubes (termed ‘Tigillites’ in the local literature) and other simple trace fossils can be seen. These beds identified as the Goulimine Formation on the Bou Izakarn map sheet are probably not coeval with those assigned to the formation on the Foum-el-Hassane map sheet. Lithologic and stratigraphic comparison suggests that the formation is partly coeval with the Jbel Afraou Formation.
December 5 STOP 17. TIMOULAYE IZDER A section ca. 2 km northeast of Timoulaye Izder exposes the upper Igoudine (Tiout Member)–middle Issafen Formations in a lithofacies succession charac-teristic of the western Anti-Atlas. The succession progressively steepens in dip from 20–27o SE on the cliff-forming Tiout
Member–Amouslek Formation, and increases to nearly vertical on the slope-forming Issafen Formation. 27 m of brownish-weathering Tiout Member are exposed on the dip slope, and are succeeded by a thick (27–312 m), deepening–shoaling succession through the
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Fig. 41. First Bani east of Timoulaye Izder, southwestern Anti-Atlas. View towards northeast shows ridge with steeple dipping sequence of Middle Cambrian rocks. Rich Khlifa Formation (poorly visible) forms lower slope to the left and is overlain by the Bailiella Formation (middle slope). The crest is formed by resistant, often well-washed quartz arenites of the Jbel Afraou Formatiom. Ordovician units form the soft lower slope to the right.
Amouslek Formation. The lowest Amouslek Formation is marked by the appearance of nodular, trilobite-bearing wackestones in green shale (27–31 m), and the lower Amouslek Formation (to 79.5 m) is dominated by massive, algal (Girvanella) buildups. Archaeocyathans may be a significant component of these build-ups (35, 39– 42, 73.5–79.5 m). The middle part of the formation (79.5– 288 m) is dominantly deeper-water, trilobite-bearing, bluish green to green siltstones and shale with thin nodular and trilobite hash limestones. Probable shoaling intervals are marked by archaeocyathan-“algal“ buildups (137–144, 163–165, 270.3–277.8, and 279.2–279.8 m) characteristically underlain by ooid-pisolite-fossil hash pack- and grainstone beds. A thin (to 10 cm-thick) volcanic ash is preserved in this deeper facies of the Amouslek Formation (105.6 m). Shallowing is obvious in the uppermost Amouslek Formation, with evidence for shallow-water, wave-dominated deposition of fine-quartz sand- and ooid/-pisolitedominated intervals (288–289, 298–300, 308.2–312 m), the appearance of gypsum molds and vugs in the top 1.8 m, and a karst at its top. Increased rates of sedimentation in the upper Amouslek Formation are suggested by balland-pillow sandstones (291.2–292.6, 296–296.6 m). This
mudstone-dominated, middle part of the Amouslek Formation is particularly fossiliferous. It features trace fossils (Planolites beverylensis, Palaeophycus, probable Cruziana), hyoliths, the characteristic obolellid brachiopod Brevipelta chouberti, trilobites (Fallotaspis appears at 79.5 m, Bigotina at 208.6 m, Daguinaspis at 279.5 m, and Resserops at 292.3 m), and the enigmatic brachiopodlike Microschedia amphitrite. The depositional sequence boundary at the top of the Amouslek is overlain by a micaceous quartz arenite (312–313.5 m) at the base of the red and green mudstonedominated Issafen Formation (312–476 m; upper part covered). The mudstones and thin nodular limestones in the Issafen represent a more distal, offshore facies, and thus, just as the siliciclastic mudstones in the underlying middle Amouslek Formation, have a higher abundance of trilobite remains than the shallower marine facies of the section (i.e., thrombolitic and oolitic facies and archaeocyathan build-ups of the Amouslek Formation). The higher abundance of trilobites in more offshore facies characterized the fossil group’s distribution from their evolutionary appearance in the late Early Cambrian through their extinction in the Permian (see WESTROP et al. 1995; LANDING & WESTROP, 2004).
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Fig. 42. Top of the Amouslek Formation and base of the Issafen Formation. Boundary is illustrated by change in the microfacies of the carbonate beds. Note cross-bedded ooid-pisolith limestones in the lower part of the photo contrasting the limonitic, dolomitized beds above. Hammer for scale. Timoulaye Izder section, southwestern Anti-Atlas.
December 5 STOP 18. AKHSAS PLATEAU The trilobite-bearing Lower Cambrian in the Anti-Atlas is generally dominated by green shales with minor intercalated limestone beds. These limestone beds are usually bioclastic and well washed or they include remains of archaeocyathan build-ups. Archaeocyathen reef complexes with considerable relief are relatively rare. A brief stop features massive limestones that totally replace the ordinary shale-dominated succession in the region around Tiznit. This limestone–dolostone complex was termed „Calcaires de Tiznit“ and probably replaces the upper Igoudine, the Amouslek and possibly the lower part of the Issafen Formation. Due to the local tectonic history and the resulting comparatively strong deformation the limestones are fairly recrystallized. Determinable faunas have not been detected yet so that
the biostratigraphic position is unclear. The massive limestones create a strange plateau (the „Plateau des Akhsas“) with karst morphology, which is unusual for Morocco. The high topographic position together with the solid rock permits the growth of the characteristic Argane trees in an unusual position close to the sea. The southern slope of the Akhsas Plateau (north of Bou Izakarn) offers a complete, spectacular section through the lower Igoudine, the Lie de vin and the Adoudou Formations. The Akhsas Plateau as a geomorphological unit is a peneplain of deeply weathered, partly karstified limestones and dolostones. This paleosurfaces lies roughly at 800–1000 m above sea-level and represents the socalled ‘pre-Hamada’ surface.
December 5 ROAD FROM TIZNIT TO AGADIR Tiznit is the gate to the deep south of Morocco and well-known for its superb silver market downtown. About 5km to the west of Tiznit, at the sea shore, is a place termed Sidi Moussa d’Aglou, where the first archaeocyathans were found in Morocco in 1926 (BOURCART 1927; BOURCART & LE VILLAIN 1928a 1928b, 1931), synchronously the first discovery of fossiliferous Lower Cambrian in Africa. The road north to Agadir passes Oued Massa close
to its mouth. The estuar of Oued Massa is seen as the place where the whale spit out the prophet Jona. The area is nowadays the largest natural preserve of southern Morocco, established as a refuge for sea birds and water fowl. The drive to Agadir via Inezgane crosses Oued Souss close to its mouth. This is the largest river in the Souss Basin and separates the High Atlas ranges from the AntiAtlas.
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Acknowledgments This article is a contribution to the Cambrian Subdivision Project of the International Subcommission on Cambrian Stratigraphy. We are indebted to W. HELDMAIER for his assistance during the preparation of field excursion notes during the MOROCCO ’95 symposium from which a number of data are used and republished herein and for his generosity in supplying additional data from the Issafen Syncline and Hassi Brahim.
We further thank E. BERNEKER and the late K. SDZUY for unpublished data. E. BERNEKER also assisted with determinations of archaeocyathans. E. L. acknowledges field and laboratory assistance from the National Science Foundation (grant EAR 94-15773). The field work of G. G. was partly financed by the Deutsche Forschungsgemeinschaft (DFG).
A contribution to the Cambrian Subdivision Project of the International Subcommission on Cambrian Stratigraphy
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