Neoproterozoic uppermost Haut-Shiloango Subgroup (West Congo Supergroup, Democratic Republic of Congo): Misinterpreted stromatolites and implications for sea-level fluctuations before the onset of the Marinoan glaciation

Journal of African Earth Sciences 90 (2014) 49–63

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Neoproterozoic uppermost Haut-Shiloango Subgroup (West Congo Supergroup, Democratic Republic of Congo): Misinterpreted stromatolites and implications for sea-level fluctuations before the onset of the Marinoan glaciation F. Delpomdor a,⇑, F. Kant b, A. Préat a a Université libre de Bruxelles, Department of Earth and Environmental Sciences, Unit of Biogeochemistry and Modeling of the Earth System, Avenue F.D. Roosevelt 50 CP-160/02, B-1050 Brussels, Belgium b Centre de Recherches Géologiques et Minières, Avenue de la Démocratie 44, Gombe/Kinshasa, The Democratic Republic of the Congo

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Article history: Received 18 January 2013 Received in revised form 12 November 2013 Accepted 14 November 2013 Available online 23 November 2013 Keywords: Haut-Shiloango Subgroup Pre-glacial carbonate deposits Lithofacies Chemostratigraphy REE+Y geochemistry

a b s t r a c t The middle Neoproterozoic carbonate-dominated uppermost Haut-Shiloango Subgroup (Sh8h and Sh8i members) in the Lower-Congo Province of the Democratic Republic of Congo is considered as recording pre-glacial shallow-marine sedimentation with stromatolitic reefs overlain by the Upper Diamictite Formation. We investigated these stromatolitic carbonates in order to highlight their biogenicity. Newly defined lithofacies and geochemical analyses (stable isotopes, major, trace and REE+Y elements) are used to provide insights into the origins of the depositional events that occurred immediately before Marinoan global glaciation. These insights should in turn provide constraints on the models developed for this glaciation event. The series consists of three shaly and carbonate lithofacies: (i) alternating limestones and claystones (lithofacies 1); (ii) nodular wackestones (lithofacies 2); and (iii) clast-supported conglomerates and breccias (lithofacies 3). Lithofacies 1 is an open marine low-energy mid/outer ramp system with hummocky cross-laminations and distal tempestites; lithofacies 2 is a distal slope facies with synsedimentary contorted structures, slided and slumped semi-consolidated limestone beds; lithofacies 3 consists of debris flows deposited in a basinal setting controlled by synsedimentary faults. None of the facies exhibits petrographic evidence of biogenicity such as stromatolitic laminar-reticulate fabrics and/or associated sediments (e.g. peloids, oncoids, ooids) or typical features such as mudcracks or clotted fabrics. The uppermost Haut-Shiloango Subgroup is made up from the stratigraphic succession of the three lithofacies and corresponds to a deepening-upward evolution from storm-influenced lithofacies in mid- and outerramp to deep-water environments, with emplacement of mass flow deposits in toe-of-slope settings. These processes occurred along tectonically active continental margins locally influenced by altitude glaciers, developed after a rift–drift transition. Uniform flat non-marine shale-normalized REE+Y patterns indicate freshwater-influenced signatures in the Sh8h carbonates. Moderate Y, Zr and Rb values reflect continental detrital inputs in nearshore environments rather than in deep-water environments. These nearshore sediments have been reworked from shallow inner- to mid-ramp settings into deeper outer-ramp and deep-water slope environments as a consequence of the tilting and uplifting of blocks. The blocks belonged to a graben-like basin related to the 750–670 Ma oceanic spreading in the central-southern Macaúbas Basin. Ó 2013 Elsevier Ltd. All rights reserved.

1. Introduction The Neoproterozoic Era heralded dramatic climatic changes, global carbon cycling and atmospheric oxygen budget (Knoll et al., 1986; Derry et al., 1992; Knoll, 1992; Des Marais, 1994; ⇑ Corresponding author. Tel.: +32 026502901. E-mail address: [email protected] (F. Delpomdor). 1464-343X/$ - see front matter Ó 2013 Elsevier Ltd. All rights reserved. http://dx.doi.org/10.1016/j.jafrearsci.2013.11.008

Strauss, 1997; Hoffman et al., 1998; Canfield, 1999) deeply intertwined with global tectonics (Kaufman et al., 1993, 1997; Halverson et al., 2007; Young, 2012). During this time, the low-latitude distribution of continents (Kirschvink, 1992; Hoffman et al., 1998; Hoffman and Schrag, 2002; Li et al., 2008) would have favoured a cool climate due to the nature of continental chemical weathering, atmospheric and oceanic circulation, and the magnitude of planetary albedo (Halverson et al., 2010). These effects led to

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several icehouse climate periods distributed at wide intervals throughout three or four global glaciations (poorly constrained c. 770–735 Ma Kaigas, well constrained c. 715–680 Ma Sturtian, well constrained c. 660–635 Marinoan, extremely well constrained c. 585–582 Ma Gaskiers; Frimmel et al., 1996; Dempster et al., 2002; Kendal et al., 2004; Lund et al., 2003; Fanning and Link, 2004; Hoffman et al., 2004; Zhou et al., 2004; Condon et al., 2005). These glaciations were followed by a sudden return to a greenhouse climate due in part to the increase of atmospheric carbon dioxide from volcanic degassing, enhancing a greenhouse warming of the surface of the Earth (Snowball Earth Hypothesis; Kirschvink, 1992; Hoffman et al., 1998; Kennedy et al., 2001; Hoffman and Schrag, 2002). The icehouse and greenhouse periods are evidenced by the presence of diamictites and cap carbonates throughout all continents (Hoffmann, 1989; Young and Gostin, 1991; Young, 1995; Wang and Li, 2003; Deynoux et al., 2006; Allen et al., 2007; Li et al., 2008). However, a number of other causative hypotheses have been proposed as alternative explanations for the Snowball Earth model (carbonate burial in Robert, 1971, 1976; high orbital obliquity in Williams, 1993, 2000, 2008; ocean stagnation in Kaufman et al., 1993; Grotzinger and Knoll, 1995; continental or zipper-rift models in Eyles, 1993; Young, 1995; Eyles and Januszczak, 2004). Although the Neoproterozoic has been widely studied using stable isotope geochemistry to establish regional or global inter-basinal correlations (Derry et al., 1989; Asmerom et al., 1991; Kaufman and Knoll, 1995; Bartley et al., 2001; Melezhik et al., 2001; Shields and Veizer, 2002; Halverson et al., 2005, 2010; Kaufman et al., 2006), the interglacial depositional environments remain poorly constrained, with a few prominent exceptions concerning the Trezona and Yaltipena formations in Australia (Preiss and Forbes, 1981; Lemon and Reid, 1998; Sohl et al., 1999; Rose et al., 2012), the Elbobreen Formation of Spitsbergen (Knoll and Swett, 1990; Fairchild and Hambrey, 1984, 1995), and the Keele Formation in Canada (Day et al., 2004). The upper carbonate Haut-Shiloango Subgroup, i.e. the Sekelolo unit, in the Democratic Republic of Congo (DRC) represents an excellent type section to study the Neoproterozoic interglacial sedimentation. The Haut-Shiloango Subgroup occurs locally several hundred metres above glacial diamictites of the Lower Diamictite Formation and immediately below glacial diamictites of the Upper Diamictite Formation, which correlate with the Sturtian and Marinoan global glaciations, respectively (Frimmel et al., 2006; Tait et al., 2011). The upper Haut-Shiloango Subgroup has previously been interpreted as recording (i) a marine carbonate transgression with in situ stromatolitic biohermal reefs (Cahen, 1950, 1954, 1978), or (ii) a post-Sturtian open marine carbonate platform succession (Frimmel et al., 2006; Frimmel, 2009, 2010). In situ preglacigenic stromatolites are therefore of particular importance since they indicate that shallow water environments preceded the glacial episode (Cahen, 1950). No study or no sampling has been carried out since 1950, although new data have been made available from coeval levels in the neighbouring areas (present day Brazil, Gabon, Namibia, Australia, Spitsbergen). Our study contributes to this gap and includes new field sedimentological observations completed by examination under a petrographic microscope. We also sutidied stable istopes, and carried out elemental geochemical analyses. The purpose of this paper is to (re)interpret the paleoenvironments and the sedimentological processes in the carbonate succession of the uppermost Haut-Shiloango Subgroup, and to investigate on their real biogenicity as previously defined. We examine the general depositional setting that occurred immediately before the Marinoan global glaciation, which in turn should provide constraints on the models developed for this glaciation event.

2. Geological setting 2.1. Regional context The West Congo Belt (WCB) is exposed for 1300 km along the western margin of the Congo Shield. It comprises a thrust-foldbelt with a top-to-(north-)east transport direction from the hinterland domain in the west and grades into a foreland domain in the east with decreasing regional metamorphism from amphibolite facies (in the west) to unmetamorphosed rocks (in the east) (Frimmel et al., 2006). Lithostratigraphically, the West Congo Supergroup (Fig. 1A) is divided, from oldest to youngest, into the rift-related volcanoclastic 999 ± 7 Ma Zadinian and volcano-sedimentary ±910 Ma Mayumbian groups deposited on a ±2.1 Ga polymetamorphic Kimezian basement, and the sedimentary West Congolian Group (Tack et al., 2001). The West Congolian Group (originally described by Cahen, 1978) is subdivided into four subgroups: (i) the 1650 m-thick continental siliciclastic and rift-related volcanic 980–920 Ma Sansikwa Subgroup (Frimmel et al., 2006; Straathof, 2011); (ii) the 700–800 m-thick siliciclastic and carbonate 650–800 Ma Haut-Shiloango Subgroup (Frimmel et al., 2006; Poidevin, 2007; Delpomdor and Préat, 2013); (iii) the 1000-m-thick carbonate 575–550 Ma Schisto-Calcaire Subgroup including cap carbonates (C1), stromatolites and oolites (Delhaye and Sluys, 1923; Bertrand-Sarfati, 1972; Frimmel et al., 2006; Poidevin, 2007; Straathof, 2011; Delpomdor and Préat, 2013); and the approximately 1000 m-thick red-bed siliciclastic 566 ± 42 Ma Mpioka Subgroup (Frimmel et al., 2006). Up to two diamictite units, i.e. the 400 m-thick Lower Diamictite and the 200 m-thick Upper Diamictite formations (Cahen, 1948, 1950; Lepersonne, 1951; Cahen and Lepersonne, 1981) have been interpreted, respectively, as Sturtian (U–Pb age: 694 ± 4 Ma; Straathof, 2011) and Marinoan in age (Frimmel et al., 2006; Tait et al., 2011). The overlying red-bed siliciclastic Inkisi Subgroup (I) is reported as Late Paleozoic, although the youngest detrital zircon grain yielded Neoproterozoic U–Pb ages of 851 ± 18 Ma (Straathof, 2011) and 558 ± 56 Ma (Frimmel et al., 2006), or Permian (Alvarez, 1995) in age.

2.2. Age of the Haut-Shiloango Subgroup (Fig. 1A) The age of the Haut-Shiloango Subgroup is poorly constrained due to a lack of geochronological data, i.e. igneous intrusions or volcanic ash deposits, and biostratigraphical fossils. The age of the Sansikwa Subgroup is even less well known, but has been inferred by Frimmel et al. (2006) to be c. 750 Ma on the basis of correlations with the overlying Lower Diamictite Formation and the Kaigas glacial event in southern Namibia. New U–Pb dating of baddeleyite grains from the dolerite Sumbi-type sill intruding the Sansikwa Subgroup and the Lower Diamictite Formation yielded a crystallization age of 694 ± 4 Ma (Straathof, 2011). In the timeequivalent Louila Formation in Gabon, rhyolitic tuffs yielded an U–Pb SHRIMP age of 713 ± 49 Ma (Thiéblemont et al., 2009) consistent with the 709 ± 20 Ma detrital zircon age of the upper clastic unit of the Haut-Shiloango (Frimmel et al., 2006). Although no specific trend can be deduced from the C isotope data of the HautShiloango Subgroup (Frimmel et al., 2006; Straathof, 2011; the present paper), it is worth noting that the limestones are characterized by anomalously high d13C ratios similar to the positive d13C excursion in post-Sturtian carbonates elsewhere (Frimmel et al., 2002; Halverson et al., 2005). Near-primary 87Sr/86Sr ratios (0.7068–0.7072; Frimmel et al., 2006; Poidevin, 2007) suggest deposition between 650–800 Ma on the basis of comparison with the composite Sr isotope curve (Jacobsen and Kaufman, 1999; Halverson et al., 2007). Detrital zircon grains from the Upper

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Fig. 1. (A) Synthetic lithostratigraphic logs of the West Congo Belt in western Central Africa (modified after Tack et al., 2001). Abbreviations:   U–Pb SHRIMP zircon ages (Tack et al., 2001);    87Sr/86Sr chemostratigraphy (Poidevin, 2007);     207Pb–206Pb detrital zircon ages (Frimmel et al., 2006);      U–Pb ages on baddeleyite grains (Straathof, 2011). (B and C) Detailed stratigraphy, lithology and sedimentary structures of the two studied outcrops in Safricas quarry crossing the Sh8h and Sh8i members of the HautShiloango Subgroup, the Upper Diamictite Formation and the cap carbonate C1 Formation of the Schisto-Calcaire Subgroup; (D) Detailed lithologs of the BCM outcrops in the Lamba Rock crossing the Sh8h member of the Haut-Shiloango Subgroup; (E) Detailed lithologs of the BCM outcrops alongthe northern part of the Congo dia Kati anticline crossing the Sh8h member of the Haut-Shiloango Subgroup.

Diamictite Formation yielded a maximum U–Pb age of 707 ± 23 Ma (Straathof, 2011) consistent with the age of the Marinoan Ghaub Formation at 635.5 ± 1.2 Ma (Hoffman et al., 2004). The HautShiloango Subgroup is therefore considered to have been deposited in a time-interval between 694 ± 4 Ma and c. 635 Ma. 3. Materials and methods 3.1. Samples The study area lies along the Congo dia Kati anticline (Fig. 2A), situated about 200 km south-west of Kinshasa in the Lower-Congo Province (DRC). The studied section crops out in a well-preserved quarry-cut and poorly-exposed cliffs section covered by dense grass savana. Samples for petrographic studies were collected from the Safricas quarry (Sh8h members to C1 Formation, including the Upper Diamictite Formation; BCJ and BCL series, Figs. 1B, C, 3A and B), from the Lamba Rock (Sh8h member; BCM series, Figs. 1D and 4A) and from near the village of Kango (Sh8h member; BCT series, Figs. 1E and 5A), successively, along the N1 highway near Kimpese. Our studied sections are mapped in Fig. 2B. Thirty-two samples of the Sh8h to Sh8i members were taken at vertical half-metrespaced intervals. Three lithofacies were distinguished according to their lithology, sedimentary structures, and bed geometry. To detail the geometry and internal structures of the beds, we prepared drawings from photographic sketches. To clarify the origin

of the deposits, we complemented the macroscopic field study with a study of thin sections. Each sample was cut into slabs and each slab polished to provide detailed pictures of sedimentary and textural features. The carbonate classification was used to give a general description of the components and matrices (Dunham, 1962; Embry and Klovan, 1972; Sibley and Gregg, 1987). To identify carbonate mineral components, we etched all thin sections with Alizarine red S and potassium ferricyanide. 3.2. Carbon and oxygen isotopes Twenty samples were selected for whole-rock carbon and oxygen isotope analyses. Carbonate powders were reacted with 100% phosphoric acid (density >1.9; Wachter and Hayes, 1985) at 75 °C using a Kiel III online carbonate preparation line connected to a ThermoFinnigan 252 mass spectrometer. All values are reported in per mil relative to V-PDB by assigning a d13C value of +1.95‰ and a d18O value of +2.20‰ to NBS19. The reproducibility of the d13CV-PDB and d18OV-PDB measurements is 0.04‰ and 0.07‰ (1r) respectively. The analyses were performed at Erlangen University in Germany (Prof. M. Joachimski). 3.3. Elemental geochemistry Major element, trace element and REE+Y concentrations were investigated on four unaltered samples, and subjected to an aqua

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Fig. 2. (A) Sketch of geological map of Kimpese area and locations of studied outcrops in the dashed rectangle (Lepersonne, 1974). (B) Detailed studied area in the Congo dia Kati anticline (after Cahen (1950)) and location of outcrops (stars).

Fig. 3. Safricas quarry. (A) Photographs of the five studied outcrops located in the eastern part of the Congo dia Kati anticline; (B) Simplified pattern of the studied area with lithostratigraphic units and located samples. Squares (A to E) are explained in Fig. 6.

regia digestion. Aqua regia consists of a 0.5 g digested sample at 90 °C in a microprocessor controlled digestion block for 2 h. The solution is diluted and analyzed by ICP/MS using a Perkin Elmer SCIEX ELAN ICP/MS at Activation Laboratories Ltd. in Canada. REE+Y concentrations were normalized with the Post-Archean Australian Shales (PAAS) composite (Taylor and McLennan, 1985). Particular attention was paid to elevated REE contents,

which could result from contamination by oxides, sulphides, phosphates or silicates, derived from hydrothermal input and/or from detrital shales. Concentrations of Zr and Al were used as further proxy indicators of shale contamination. Shale-normalized element anomalies calculated between neighbouring pairs are constant: La/La⁄ = La/(3Pr–2Nd), Ce/Ce⁄ = Ce/(2Pr–Nd), Eu/Eu⁄ = Eu/ 0.67Sm + 0.33 Tb), and Gd/Gd⁄ = Gd/(2 Tb–Dy). Anomalies were

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Fig. 4. Lamba Rock. (A) Photograph of the Lamba Rock with location of Fig. 3; (B) Mud-supported autoclastic brecciated horizons around storm-influenced plastoclasts; (C) Contorted thin-bedded limestones with millimetre-scale laminations of pale red to rust-coloured calcareous mudstones.

Fig. 5. Northern part of the Congo dia Kati anticline. (A) Photograph of the BCM series outcrop; (B) Gently contorted limestone beds with millimetre-scale laminations of pale red to rust-coloured calcareous mudstones.

also calculated from a geometric average, assuming that the ratio between near neighbour concentrations is constant as follows (Lawrence et al., 2006): La⁄ = Pr⁄(Pr/Nd)2, Ce⁄ = Pr2⁄Nd, Eu⁄ = (Sm2⁄Tb)0.33 and Gd⁄ = (Tb2⁄Sm)0.33. Differences in results obtained by these two methods are around 5%. The very low REE concentrations necessitating sophisticated analytical techniques and instrumentation for measurement are not given because the detection measurement is limited to <0.1. We estimated that in the HautShiloango Subgroup the total REE concentrations were high (>25 ppm). 4. Uppermost Haut-Shiloango Subgroup – Sekelolo unit The Haut-Shiloango Subgroup (Sh) consists of two units – Mouyonzi and Sekelolo (Lepersonne, 1951, 1974). The lower two-thirds of the Mouyonzi unit (ShI unit) are composed of 450–650 m-thick siliciclastics divided into five formations (Sh1 to Sh5). The overlying Sekelolo unit (ShII unit) consists of a 200–250 m-thick siliciclastic–carbonate series, composed of 40–50 m-thick creamcoloured coarse-grained feldspathic quartzites (Sh6 formation), 20–63 m-thick grey-greenish shales (Sh7 formation), ±180 m-thick blue-greyish limestones, locally clayey, alternating with dark shales (Sh8 formation). The Sh8 formation is subdivided into nine members, i.e. Sh8a to Sh8i (Cahen, 1950). In this study, we re-examined the Sh8h and Sh8i members. Unlike previous lithostratigraphic descriptions, our observations showed that the breccia-like facies of the Sh8i member is not specific to this member,

but instead varies, at intervals through Sh8h to Sh8i members, from mudstone/limestone beds to plurimetric (1–3 m-thick) lithosequences of breccias. Lateral variations were not observed due to the lack of outcrops. We summarize the lithofacies sequences in the Sh8h and Sh8i members below.

4.1. Sh8h member The Sh8h member consists of a ±20 m-thick carbonate succession of (from base to top) gently contorted thin-bedded blue-grey fine-grained limestones evolving to disrupted and contorted slumped blocks of blue-grey limestones re-oriented sub-parallel to bedding, and toward the top to clast-supported breccias. Sedimentary structures in the bedded limestones include parallel and convolute laminations, ‘hummocky’ cross-laminations and ripples. Rotated slumped blocks of limestone and slide scars are also observed. The laminated facies is made up of contorted laminae, less than 1 cm-thick, of pale red to rust-coloured clayey carbonate materials. These laminae occur as a matrix between the large clasts from the disrupted laminated limestones. The clasts accumulated as elongated lenses sub-parallel to bedding. Breccias display 0.30–1.50 m-thick clast-supported beds with cobbles to slabs of up to 0.3–0.5 m in length, derived from the underlying Haut-Shiloango members. The lower contact between the breccias and the slumped beds is erosional or faulted with intraformational truncation surfaces.

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4.2. Sh8i member The Sh8i member consists dominantly of over 1.5–2 m of clastsupported breccias with slided huge blocks of contorted thin-bedded-derived blue-grey limestones and elongated boulders of underlying limestones. Clasts range in size from cobbles to slabs up to 0.3–0.5 m in length. The matrix consists of grey-pale dololutites. The brecciated beds are interbedded with gently contorted thin-bedded blue-grey limestones evolving to disrupted and contorted slumped blocks of blue-grey limestones re-oriented subparallel to bedding. Intraformational truncation surfaces, draping the thin-bedded limestones, occur above the bedded sequence of redeposited limestones. 5. Lithofacies 5.1. Lithofacies 1 (LF1): alternating limestones and claystones/fairweather mid/outer-ramp facies with distal mud tempestites 5.1.1. Description Blue-grey fine-grained limestones (centimetre- to decimetrethick) alternating with millimetre-scale laminae of pale red to rust-coloured calcareous mudstones (Fig. 6A). Small and medium-scale erosional structures such as gutter casts with erosive basal surfaces and rip-up clasts are preserved on the lower bedding planes of the limestones. Bed tops are often truncated or missing. Within beds, limestones show small-scale continuous or discontinuous parallel laminae overlain by hummocky cross-laminations. In thin sections, truncated tempestite sequences with erosive bases disrupting underlying facies are overlain by graded micro-layers of laminated wackestones ending with dark mudstones. The thin muddy beds are relatively planar with microstylolitic features. The contorted thin-bedded limestones pass to disrupted slide scars and rotated slumped blocks re-oriented sub-parallel to bedding

(Figs. 4C, 5B and 6D). Creep structures are associated with bed slidings along shear zones. Intraformational truncation surfaces are observed along the contact between limestones and redeposited bedded sequences (see lithofacies 2 and 3). 5.1.2. Interpretation This facies represents a low-energy outer ramp environment near or below the storm wave base. Samples show typical sequences with truncated distal mud tempestites. These sequences are mud-dominated with gently and low-angle curved cross-laminae in the wackestones and evolve to thin planar laminated mudstones. The contact with the base is commonly planar, but may exhibit erosive surfaces with gutter casts overlain by coarsegrained storm-derived intraclasts. This lithofacies is attributed to a mid- to deeper outer ramp setting relatively similar to the depositional environment of Spanish and Hungarian Muschelkalk facies (Calvet and Tucker, 1988; Török, 1998). 5.2. Lithofacies 2 (LF2): nodular mudstones-wackestones/deeper outer ramp facies 5.2.1. Description Nodular wackestones with pronounced nodule-supported fabrics of disrupted and contorted carbonate tempestites embedded in microstylolitic seams of pale red to rust-coloured calcareous mudstones (Fig. 6B and C). The nodules consist of plastoclasts ranging in size from cobbles to slabs of up to 30–50 cm in diameter and smaller nodules (up to 5 cm). The weathered surface of the plastoclasts shows small-scale planar- and hummocky cross-laminations. Several levels exhibit mud-supported autoclastic brecciated beds around tempestite plastoclasts (Fig. 4B). Each clast of breccia is surrounded by disturbed pale red to rust-coloured calcareous mudstone seams or microstylolites.

Fig. 6. Lithofacies of the Safricas quarry: (A) Gently contorted and slumped thin-bedded limestones with millimetre-scale laminations of pale red to rust-coloured calcareous mudstones; (B) Nodular wackestones surrounded by ‘microstylolitic seams’ of pale red to rust-coloured calcareous mudstones overlying disturbed limestone beds; (C) Nodule-supported fabric of disrupted and contorted tempestite-dominant beds of limestones in pale red to rust-coloured calcareous mudstones; (D) Disrupted slide scars and rotated slumped blocks re-oriented sub-parallel to bedding; (E) Clast-supported breccia with fragments or plastoclasts of underlying limestone beds .

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5.2.2. Interpretation The nodular lithofacies is classically attributed to a low-energy deeper outer ramp deposit. Nevertheless, this lithofacies is filled with coarse debris of parallel-bedded limestone slumped blocks and slides (similar to facies F of Mutti and Ricci Lucchi, 1978) indicating a carbonate slope environment. Limestones with tempestite sequences were derived from a peri-platform margin. Pressure solution processes are indicated by microstylolites around the contacts along the nodules. These microstylolites result from the reworking of small-scale plastic creeped semi-lithified beds. This mechanism could be either attributed either to bottom currents (storm induced) or to a bedload increased by sediment covers (Török, 1998). The small size of clasts (cm) and bed thickness (m) of autoclastic breccias are indicative of gravity-driven smallscale sediment flow down a very gentle slope (<1°) or along a flat ramp (Török, 1998). 5.3. Lithofacies 3 (LF3): clast-supported breccias/‘slope’ facies 5.3.1. Description The clast-supported breccias consist of normally graded subangular to subrounded cobbles to granules in a sandy to silty matrix (Fig. 6E). The breccias contain fragments of underlying limestones. In thin section, a floating grained-texture of quartz and mud-carbonate pebbles with interparticle calcite cements is observed, with a dominant grain-size around 500 lm. 5.3.2. Interpretation Huge blocks of thin-bedded slope limestones slid downslope and formed breccias (facies F; Mutti and Ricci Lucchi, 1978) transported and deposited by high-density turbidity currents and grain flows. The graded laminae record turbidity currents, which flowed downslope due to gravity (Rhodri Johns, 1978). Laminae of coarseand fine-grained facies, without erosive bases, may be the products of accelerating and decelerating turbidity current flows. Largescale slope failures revealed the variable sliding of semi-consolidated sediments along discrete basal shear planes, a process that created classical intraformational truncation surfaces in the upper slope carbonate sequences. 5.4. Lithofacies synthesis and ‘the problem of the stromatolites’ Our thin section study revealed that the facies are related to strong reworking processes that occurred as a result of storm and turbidity currents on a slope located in the distal part of a shelf (ramp setting). We observed no diagnostic stromatolite features under the petrographic microscope. If shallow-water biohermal accretion had occurred, we would expect stromatolite frameworks dominated by erected columnar microbialites made almost exclusively by laminar-reticulate fabrics, as they are well represented in the Neoproterozoic (Turner et al., 2000). However, such fabrics, consisting of a microscopic calcimicrobial texture with diffuse alternating laminae of felted prostrate calcimicrobial filaments and loosely packed oriented filaments (Turner et al., 1993), are absent in the succession studied. We observed no growth cavities, internal sediments, clotted fabrics or synsedimentary cements. Moreover, we observed no cross-stratified oncoid or oolitic limestones (packstones, grainstones) or mud-cracked facies (mudstones, wackestones) typically associated with shallow subtidal stromatolites. Neither did we observe in our study any typical features of deeper-water stromatolites (e.g. grumous micrite, renalcid encrustations, also laminar-reticulate fabrics) have been observed in our study. The lateral extent of the lithofacies in each member (Sh8h and Sh8i) on a large scale in the field (tens to hundreds of metres) did not show the kind of typically attributable to biohermal accretion.

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6. Geochemical data 6.1. Carbon and oxygen isotopes Stable isotope data from the Sh8h and Sh8i members exhibit a precipitous decline in d13C at the top of the Sekelolo unit and at the base of the Schisto-Calcaire Subgroup. The datasets are given in Table 1. In the Haut-Shiloango Subgroup, the d13C values range between +2.4‰ and +8.0‰, averaging +5.6‰ (n = 20) and the d18O range between 11.2‰ and 5.5‰, averaging 9.7‰ (n = 20) (Fig. 7). In the Sh8h and Sh8i members from Safricas quarry samples, the d13C and d18O values range between +8.0‰ and +2.4‰, and between 11.2‰ and 5.5‰, respectively. The d13C and d18O values exhibit a strongly positive covariance (mean covariance = 0.73) (Fig. 7). In the Lamba Rock samples, the d13C and d18O values are similar to those found in the Sh8h member from the Safricas quarry. Samples yielded mean d13C and d18O values of +7.6‰ (n = 2) and 8.4‰ (n = 2), respectively. The range of our values is consistent with those obtained by Frimmel et al. (2006) and Straathof (2011), who report mean d13C values of +5.5‰ (n = 10) and +5.4‰ (n = 50) respectively, and mean d18O values of 9.9‰ (n = 10) and 9.3‰ (n = 50), respectively. 6.2. Major, trace and REE+Y concentrations 6.2.1. Major and trace concentrations Our samples come from planar-bedded facies consisting of centimetre- to decimetre-thick blue-grey fine-grained limestones alternating with millimetre-scale laminations of pale red to rustcoloured calcareous mudstones. The blue-grey fine-grained limestones yielded a CaCO3 content of between 64% and 74% (n = 4) indicating a notable clay content. Analyses of whole-rock CaO content (mean CaO = 39.4%; n = 4) and MgO content (mean MgO = 0.8%; n = 4) contents indicate that the Sh8h member is a predominantly non-dolomitized limestone. Nevertheless, the Sr concentrations of between 564 ppm and 754 ppm (n = 4) seen in our samples point to former sulphate casts replaced by calcite. The high Sr concentrations are coincident with low Al2O3 concentrations (Table 1). The Zr and Th concentrations are relatively high with ranges of between 3.5 ppm and 9.8 ppm (n = 4), and 0.6 ppm and 2.4 ppm (n = 4), respectively. 6.2.2. REE+Y concentrations Samples from the Sh8h member display relatively flat REE+Y distributions (Fig. 8) and very low total REE contents (maximum REEtotal = 57.9 ± 0.1; n = 4). The datasets are given in Table 2. The REE+Y distributions reveal two different trends: (i) a relatively slight enrichment in light and middle REE (mean (Y/ Ho)PAAS = 1.0 ± 0.1 and (Nd/Dy)PAAS = 1.1 ± 0.2) in the lower Sh8h member (n = 2); and (ii) a relatively slight enrichment in light REE (mean (Y/Ho)PAAS = 1.1 ± 0.1) and a slight depletion in middle REE (mean (Nd/Yb)PAAS = 0.7 ± 0.2) in the upper part of the Sh8h member (n = 2). Both the lower and the upper parts of the Sh8h member display very low average (Y/Ho) ratios of 28.7 ± 0.1 (n = 4). These samples also show a slight positive Eu anomaly (mean (Eu/Eu⁄)PAAS = 1.31 ± 0.05), a weak positive Ce anomaly (mean (Ce/Ce⁄)PAAS = 0.98 ± 0.1), a strong positive Gd anomaly (mean (Gd/Gd⁄)PAAS = 1.3 ± 0.1) and no La anomaly. 7. Interpretation and discussion 7.1. Depositional model The assemblage of lithofacies of the uppermost Haut-Shiloango Subgroup reflects a distally-steepened ramp with mid/outer-ramp,

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F. Delpomdor et al. / Journal of African Earth Sciences 90 (2014) 49–63

Table 1 d13C and d18O values, major and trace element compositions of bulk samples from Sh8h and Sh8i members. Lamba rock Samples Member Lithology d13C (‰, V-PDB) d18O (‰, V-PDB) CaO (wt%) MgO (wt%) Al2O3 (wt%) K2O (wt%) MnO (wt%) Fe2O3 (wt%) Mn (ppm) Fe (%) Zr (ppm) Sr/Ca Mn/Sr Rb/Sr

Safricas quarry

BCM4 Sh8h lst. 7.5

BCM1 Sh8h lst. 7.7

8.1

8.7

BCL15 Sh8i lst. 2.4

BCL14 Sh8i lst. 3.0

BCL11 Sh8h lst. 4.6

10.6

10.7

10.7

BCL8 Sh8h lst. 6.6

BCL7 Sh8h lst. 7.6

BCL6 Sh8h lst. 7.3

BCL5 Sh8h lst. 8.0

BCL4 Sh8h lst. 7.6

9.6

9.9

8.8

9.7

10.3

BCL1 Sh8h lst. 7.2

BCJ29 Sh8h lst. 5.5

10.0

10.3

BCJ28 Sh8h lst. 5.7

BCJ27 Sh8h lst. 5.3

BCJ25 Sh8h lst. 4.5

10.3

10.4

5.5

BCJ24 Sh8h lst. 5.7

BCJ22 Sh8h lst. 4.8

BCJ21 Sh8h lst. 3.0

BCJ20 Sh8h lst. 4.3

BCJ19 Sh8i lst. 3.5

10.8

11.0

10.9

11.2

7.1

– – –

– – –

– – –

39.6 0.8 1.1

– – –

– – –

– – –

– – –

– – –

– – –

39.9 0.6 0.9

– – –

36.5 0.9 2.0

– – –

– – –

– – –

– – –

– – –

47.7 0.6 0.6

– – –

– – –

– – –

– – –

0.2 0.03 1.2

– – –

– – –

– – –

– – –

– – –

– – –

0.2 0.02 0.9

– – –

0.6 0.03 2.7

– – –

– – –

– – –

– – –

– – –

0.1 0.02 0.7

– – –

– – – – – –

– – – – – –

– – – – – –

196 0.8 8.1 0.0020 0.34 0.01

– – – – – –

– – – – – –

– – – – – –

– – – – – –

– – – – – –

– – – – – –

179 0.6 4.9 0.0027 0.24 0.01

– – – – – –

211 1.9 36.0 0.0021 0.38 0.04

– – – – – –

– – – – – –

– – – – – –

– – – – – –

– – – – – –

126 0.5 3.0 0.0025 0.17 0.01

– – – – – –

Fig. 8. Relatively flat shale-normalized REE+Y patterns with no distinct element anomalies in Sh8h member. Our data are consistent with those of Frimmel (2009) (dashed lines).

Fig. 7. d13C vs. d18O cross plot of carbonates of the uppermost Haut-Shiloango Subgroup (data from Frimmel et al. (2006) and Straathof (2011); and present study). Straight lines represent the covariance between d13C and d18O. Note the strong positive covariance (R2 = 0.73) at the top of the Haut-Shiloango Subgroup.

slope or base-of-slope settings. Ramp models consist traditionally of a gentle slope, commonly much less than 1°, with outer, mid and inner parts. The mid-ramp is bracketed offshore by the storm wave base and inshore by the fair-weather wave base (Burchette and Wright, 1992). The outer-ramp is the zone below the normal storm wave base, with mudstone deposition and few storm beds (Burchette and Wright, 1992). The distally-steepened ramp is related to an offshore slope break between the shallow ramp and the basin. However, the ramp may be composed of slope aprons or base-of-slope aprons adjacent to shallow-water platform margins largely composed of shallow-water materials (Pedley, 1992). The difficulty in the field of distinguishing the slope or the slope apron deposits is related to the extent of the outcrop, which does not lend itself to a three-dimensional reconstruction. The abun-

dance of slides and slumps, mud-supported breccias and intraformational truncation surfaces are however indicative of slope facies. Mullins (1983) reported modern submarine slides from the Bahamas adjacent to apron facies. Moreover, intraformational truncation surfaces result from large-scale slope failures (Cook and Enos, 1977) commonly attributed to upper slope carbonate features (Mullins and Cook, 1986). The deposition of the uppermost Haut-Shiloango Subgroup records mid/outer ramp settings, locally distally-steepened or sloped, as indicated by (i) numerous open-marine synsedimentary features (hummocky cross-laminations and distal tempestites); (ii) synsedimentary contorted structures, sliding and slumping of semi-consolidated limestone beds; (iii) synsedimentary faults (intraformational truncation surfaces); and (iv) autoclastic mudstone breccias of ‘debris flow’ origin.

7.2. Diagenetic alteration Negative d13C values, especially coupled to negative d18O values (Gross and Tracey, 1966; Banner and Hanson, 1990) are useful for deciphering diagenetic processes (Kaufman et al., 1991). A cross

57

F. Delpomdor et al. / Journal of African Earth Sciences 90 (2014) 49–63 Table 2 REE+Y geochemical data of bulk samples from Sh8h and Sh8i members. Lamba rock Samples Member Lithology La (ppm) Ce (ppm) Pr (ppm) Nd (ppm) Sm (ppm) Eu (ppm) Gd (ppm) Tb (ppm) Dy (ppm) Y (ppm) Ho (ppm) Er (ppm) Tm (ppm) Yb (ppm) Lu (ppm) REE total (ppm) Eu/Eu⁄ La/La⁄ Ce/Ce⁄ Gd/Gd⁄ Y/Ho (Y/Ho)PAAS (Nd/Yb)PAAS

Safricas quarry

BCM4 Sh8h lst. – – – – – – – – – – – – – – – –

BCM1 Sh8h lst. – – – – – – – – – – – – – – – –

BCL15 Sh8i lst. – – – – – – – – – – – – – – – –

BCL14 Sh8i lst. 5.6 12 1.4 5.5 1.1 0.3 1.1 0.2 0.9 6.1 0.2 0.5 <0.1 0.4 <0.1 35

BCL11 Sh8h lst. – – – – – – – – – – – – – – – –

BCL8 Sh8h lst. – – – – – – – – – – – – – – – –

BCL7 Sh8h lst. – – – – – – – – – – – – – – – –

BCL6 Sh8h lst. – – – – – – – – – – – – – – – –

BCL5 Sh8h lst. – – – – – – – – – – – – – – – –

BCL4 Sh8h lst. – – – – – – – – – – – – – – – –

BCL1 Sh8h lst. 5.6 11 1.3 4.6 0.9 0.2 0.9 0.1 0.8 5 0.2 0.5 <0.1 0.4 <0.1 32

BCJ29 Sh8h lst. – – – – – – – – – – – – – – – –

BCJ28 Sh8h lst. 9.3 19 2.3 8.6 1.8 0.4 1.8 0.3 1.7 9.8 0.4 1.1 0.2 1.2 0.3 58

BCJ27 Sh8h lst. – – – – – – – – – – – – – – – –

BCJ25 Sh8h lst. – – – – – – – – – – – – – – – –

BCJ24 Sh8h lst. – – – – – – – – – – – – – – – –

BCJ22 Sh8h lst. – – – – – – – – – – – – – – – –

BCJ21 Sh8h lst. – – – – – – – – – – – – – – – –

BCJ20 Sh8h lst. 2.7 5.9 0.7 2.7 0.5 0.1 0.6 <0.1 0.5 3.5 0.1 0.3 <0.1 0.3 0.1 18

BCJ19 Sh8i lst. – – – – – – – – – – – – – – – –

– – – – – – –

– – – – – – –

– – – – – – –

1.1 1 0.7 31 1.1 1.2

– – – – – – –

– – – – – – –

– – – – – – –

– – – – – – –

– – – – – – –

– – – – – – –

1.2 1 0.9 2.5 25 0.9 1

– – – – – – –

0.9 1 1 1 25 0.9 0.6

– – – – – – –

– – – – – – –

– – – – – – –

– – – – – – –

– – – – – – –

1.8 1 1 1.1 35 1.3 0.8

– – – – – – –

plot of d13C vs. d18O in the uppermost Haut-Shiloango Subgroup (Fig. 7) shows two distinct trends with a quite good covariation within the Sh8h and Sh8i members: (i) R2 = 0.82 (n = 9) for the BCL series; (ii) R2 = 0.66 (n = 9) for the BCJ series; and (iii) no analyses of the BCM and BCT series were carried out due to a lack of data. Such a high covariation points to a mixing of carbonate sediments with isotopically light meteoric fluids during early diagenesis (Delpomdor and Préat, 2013). Possible diagenetic processes causing a negative shift in carbon and oxygen isotopes are 13Cand 18O-depleted fluids related to decarboxylation during deep burial of carbonates and organic matter, early organic diagenesis or meteoric alteration during carbonate stabilization (Le Guerroué and Cozzi, 2010). Sr/Ca elemental ratios and Mn concentrations also provide a sensitive indication of diagenesis, because during exchange with diagenetic fluids, Mn is commonly incorporated into carbonates, while Sr is flushed from the carbonate lattice (Brand and Veizer, 1980; Veizer, 1983; Kah, 2000). Compared to modern seawater Sr/Ca ratios, estimated at 8.6  10 3, the Sh8h and Sh8i samples have low Sr/Ca ratios of between 2.0  10 3 and 2.7  10 3 (Fig. 9A). Kah (2000) suggests that falls of Sr/Ca ratios and increases in Mn concentrations in carbonates are related to mixing with meteoric waters, and that increases of Sr/Ca ratios are related to seawater-derived evaporitic brines. High Mn concentrations may also be related to meteoric and/or burial diagenesis (Kaufman and Knoll, 1995; Jacobsen and Kaufman, 1999; Derry, 2010). In addition, variations in Mn, like those in Fe, record similar trends for the palaeobathymetry of our reconstituted sequences. The high Mn and Fe concentrations in the Sh8h and Sh8i samples (Fig. 9B), compared to other time-equivalent post-Marinoan carbonates, suggest that our carbonates were deposited in oxic to suboxic conditions, caused by a sea-level transgression with Mn and Fe inputs derived from continental weathering or mid oceanic ridges. Note that in the last case, ophiolites, or equivalent oceanic material, have never been reported as occurring in the region during the time interval we are considering.

Our results show a d18O decrease followed by a Mn increase indicative of meteoric or mixed-water diagenetic alterations (Fig. 9C). Kah (2000) reported that the dolomitization stabilization through fresh, meteoric or mixed-water diagenesis should result in a rapid decrease in oxygen isotopic composition followed by a Mn increase as diagenesis evolved, providing reducing conditions that would facilitate the incorporation of the Mn. Nevertheless, the Sh8h and Sh8i member carbonates have low Mn concentrations (average 178 ppm; n = 4), similar to those of most Neoproterozoic carbonates (<1000 ppm) with larger local variations than those recorded in cap carbonates (Yoshioka et al., 2003; Frimmel, 2009; Zhao et al., 2009). As it is commonly accepted that during marine as well as meteoric diagenesis, Mn concentrations increase while those of Sr decrease, the Mn/Sr ratio is a good indicator of the degree of alteration (Jacobsen and Kaufman, 1999). Bartley et al. (2001) estimated that non-diagenetic carbonate rocks are characterized by Mn/Sr = 1.5 and Rb/Sr = 0.01 in limestones, and Mn/Sr = 3.0 and Rb/Sr = 0.01 in dolostones. We considered these ratios as good indicators of the degree of diagenetic alteration. A cross plot of Mn/Sr vs. Rb/Sr (Fig. 9D) shows slightly weak-altered primary signatures, and that two samples retained a primary signature, suggesting that the Sh8h and Sh8i member carbonates did not undergo diagenesis. 7.3. Marine or non-marine depositional environment According to Frimmel (2009, 2010), the calcipelitic and calcareous Haut-Shiloango Subgroup showed non-marine REE+Y signatures; this is also our conclusion based on our geochemical results of uniform flat shale-normalized REE+Y patterns without elemental anomalies. Furthermore, the freshwater component displays uniform unfractioned REE+Y patterns with only mild light and heavy REE depletions (Lawrence et al., 2006). The river water influenced PAAS-normalized REE+Y patterns show typical modest LREE enrichment, absence of seawater-like Ce and La anomalies, low Y/Ho ratios and high Al concentrations. By contrast, the normal

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F. Delpomdor et al. / Journal of African Earth Sciences 90 (2014) 49–63

Fig. 9. Diagenetic alteration. (A) Sr/Ca vs. Mn elemental concentrations. All Sh8h and Sh8i carbonates have low Sr/Ca ratios and moderately elevated Mn concentrations. They record non-marine fluid (fresh, meteoric, mixed-water or evaporitic brines) influences; (B) Mn vs. Fe concentrations; (C) d18O vs. Mn cross plot indicating possible meteoric or mixed-water diagenetic alterations; (D) Mn/Sr vs. Rb/Sr ratios. Non-diagenetic carbonate rocks have Mn/Sr < 1.5 and Rb/Sr < 0.01. Data from Frimmel et al. (2006) with dashed lines. (See above-mentioned references for further information.)

seawater REE+Y distribution is characterized by a typical uniform light REE depletion, enrichment in La, depletion in Ce, slight enrichment in Gd and positive Y anomaly (Zhang and Nozaki, 1996). The relatively unfractioned REE+Y patterns of the Sh8h member instead show a typically modern river water pattern (Lawrence et al., 2006) and/or a non-marine carbonate pattern (Bolhar and Van Kranendonk, 2007). Frimmel (2010) suggested that shale contamination had possibly influenced the overall REE+Y patterns. Concentrations of Th, Al and Zr are commonly used to assess shale contamination, as they are present in detrital material containing clay minerals (Caron et al., 2010). Moreover, low

SiO2 (0.4–3.1 wt%), Y (62 ppm), Zr (616 ppm) and low Rb (65 ppm) contents reflect minimal continental detrital input, consistent with a distal environment (Frimmel, 2009). By contrast, higher Y, Zr, and Rb concentrations are indicative of larger continental detrital inputs in a proximal, nearshore environment (Frimmel, 2009). The Sh8h member exhibits low Y and Rb concentrations with values ranging, respectively, between 3.5 ppm and 9.8 ppm (n = 4), and between 5.6 ppm and 7.8 ppm, except for one sample with 21.6 ppm of Rb. Low Zr concentrations in the Sh8h member range between 3.0 ppm and 8.1 ppm, except for one sample with 36 ppm (n = 4). A cross plot of Zr

Fig. 10. (A) Zr vs. (Nd/Yb)PAAS (black circles) and Al2O3 vs. (Nd/Yb)PAAS (white circles) cross plots; (B) (La/La⁄)PAAS vs. Zr (black circles) cross plot. Data from Frimmel et al. (2006) with dashed lines.

F. Delpomdor et al. / Journal of African Earth Sciences 90 (2014) 49–63

concentrations vs. (Nd/Yb)PAAS (Fig. 10A) shows a very low correlation (R2 = 0.0943; n = 3) between depletion or enrichment of LREE and elevated Zr concentrations. Al concentrations are also used as a proxy for shale contaminations. Most samples have low Al concentrations compared to the average for siliciclastic-contaminated carbonates of 0.42%, as calculated by Veizer (1983), and show a weak correlation with the LREE (R2 = 0.2625; n = 4; Fig. 10A), confirming possible shale contamination. Furthermore, a cross plot of (La/La⁄)PAAS vs. Zr cross plot (Fig. 10B) indicates contamination (mean Zr concentrations = 5.3%) and a distinct positive La anomaly. Calculated Y/Ho ratios range between 24.5 and 35.0, which are very different from Y/Ho ratios in open seawater, which range between 80 and 150, and are strongly dependent on the salinity (Nozaki et al., 1997; Lawrence et al., 2006). The lower Y/Ho ratios occurring in nearshore environments are estimated to be between 65 and 70 (Nozaki et al., 1997; Meyer et al., 2012). In comparison, the Y/Ho ratio in river water is equal to, or slightly higher than, the PAAS values and below seawater values, i.e. below 60 (Lawrence et al., 2006). Low Y/Ho values would therefore seem to indicate a freshwater influence during the Sh8h carbonate deposition. Moreover, the cross plot of (Y/Ho)PAAS vs. (Nd/Yb)PAAS (Fig. 11) does not exhibit any correlation suggesting a weak REE depletion and enrichment compared with (Y/Ho)PAAS further indicating a possible freshwater influence. Lozenge- and rhomboedral-shaped pseudomorphs, most probably after gypsum were observed in the thin sections. Similar recrystallization in the Upper Anisian-Middle Ladinian mid-ramp carbonates from Hungarian Muschelkalk, similar recrystallization has been interpreted as burial diagenetic effects in a dolomitized ‘slope’ related to an initial steepening of the ramp (Török, 1998). Furthermore, Sr concentrations of between 554 ppm 754 ppm, moderate when compared with those obtained by Azmy et al. (2001) in Vazante carbonates (Brazil), and high d13C values also suggest a dolomitization. However, low to moderate Sr concentrations combined with low Sr/Ca ratios reveal a possible freshwater alteration, with a schizohaline dolomitization related to freshwater Sr-depleted fluids (Pierson, 1981) or that initial Sr/Ca ratios of carbonates were low. Frimmel (2010) indicated that high d13C values may result from carbonate deposition in a nearshore setting where the water chemistry did not correspond to that of the coeval open ocean. A possible explanation for low d13C values is that they result from mixing caused by biological pumping (Shen et al., 2005) of different proportions of an isotopically light carbon reservoir from deeper to shallower water by upwelling processes in anoxic or suboxic depositional environments, where the isotopic gradient remained unchanged due to the sluggish vertical circulation of the basin. The consequences of biological pumping caused (i) a prefer-

Fig. 11. (Y/Ho)PAAS vs. (Nd/Yb)PAAS cross plot. Dashed lines are indicate of data from Frimmel et al. (2006).

59

ential uptake of 12C by primary producers followed by downward transport and remineralisation of 13C-depleted carbonates or organic matter in deep waters; and (ii) an increasing downward flux of organic matter where the surface waters become more enriched in 13C (Shen et al., 2005). In contrast, remineralization of organic matter in deep water environments releases light carbon that could be transported to the photic zone through upwelling and homogenize the carbon chemistry of the ocean (Kump, 1991, sensu Shen et al., 2005). In conclusion, the uniform flat non-marine shale-normalized REE+Y patterns in the Sh8h carbonates constitute a record of a freshwater influence, variably contaminated by detrital material. According to Banner et al. (1988), who reported little apparent effect on REE signatures during dolomitization in Carboniferous carbonates, the initial seawater-like REE+Y signature of the Sh8h member has not been significantly modified by dolomitization processes. The relatively moderate Y, Zr and Rb contents reflect continental detrital inputs in nearshore environments rather than in distal deep-water environments. In terms of palaeoenvironmental reconstruction, the non-marine REE+Y patterns combined with major and trace element data reflect signatures of non-marine sediment derived from nearshore environments towards the distally outer-ramp and slope facies. 7.4. Regional implication The western margin of the Congo Shield was affected in Early Neoproterozoic times by a rifting related to the break-up of Rodinia, c. 1.0 Ga, and the opening of the Adamastor Ocean (Tack et al., 2001) followed by passive margin-type sedimentation. Alvarez and Maurin (1991) recognized four major tectono-sedimentary episodes in this Group: (i) an infilling of rift systems with the Bouenza and the Niari Tillite formations; (ii) a eustatic cycle of the Schisto-Calcaire Group; (iii) a spreading of the Pan-African molasses of the Mpioka Group; and (iv) an establishment of deltaic sedimentation of the Inkisi (Sub)Group. Structural data (Gioan et al., 1989) and geophysical data (Godivier et al., 1986) showed that the rift initiated between 950 Ma and 700 Ma along NE–SW fault systems (Alvarez and Maurin, 1991). In the Araçuai Belt, 816-Maold oceanic crust has been reported and accompanied the extensional regime during the early Neoproterozoic continental rifting of Rodinia. This rifting took place at similar times in other areas as recorded in the Pan-African domains of southern Africa: around 795 Ma in the Zambezi Belt (Hargrove et al., 2003) and between 771 (±6) and approximately 740 Ma in the Gariep Belt (Frimmel et al., 2001, 2002), and between 1000 Ma and 700 Ma in the WCB (Tack et al., 2001; Frimmel et al., 2006; Straathof, 2011). The rifting episode caused the formation of several faulted basins in the western part of the Congo Shield, which was downwarped beneath NE–SW trending fault basins known as the Lower-Congo aulacogen in DRC (Schroeder, 1981), the Sangha (Poidevin, 1985; Vicat et al., 1989) and the Comba aulacogens in the Republic of Congo or ‘RC’ (Alvarez and Maurin, 1991). The southern extremity of these aulacogens is connected with the NW–SE orthogonal Nyanga-Niari Basin which is composed of a WCB metamorphosed western foreland or hinterland domain, and an eastern foreland domain of weakly metamorphosed and unmetamorphosed rocks. As Neoproterozoic deposits to the north and south of the Lower-Congo aulacogen can be assumed to be coeval, the Louila and Bouenza formations in Gabon are coeval with the Haut-Shiloango Subgroup in DRC and Angola (Chevallier et al., 2002). The Haut-Shiloango Subgroup deposition took place after the rift–drift transition on a passive continental margin (Frimmel et al., 2006). Wagner and Wilhelm (1971) recognized two foldings in the Louila series from Kouilou-Niari (RC), coeval with the Haut-Shiloango Subgroup (DRC). In the Lower-Congo Province, Cahen

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(1978) reported several tectonic phases with an early phase in the Haut-Shiloango Subgroup and the late ones in the Mpioka Subgroup. Note that he did not mention tectonic phases in the Schisto-Calcaire Subgroup. Recently, Pedrosa-Soares and Alkmin (2011) proposed a pre-collisional 750–670 Ma ‘E6 Event’ in the central-southern Macaúbas Basin. In terms of regional tectonic settings, lateral variations of facies and thickness showed that the subsidence was more pronounced in the Haut-Shiloango Subgroup than in the Sansikwa Subgroup (Cahen, 1978). In the Lower-Congo aulacogen, the tectonic activity was probably related to the ‘E6 Event’ approximately 750–670 Ma. This event manifested as a reactivation of former faults with synsedimentary differential tilting and uplifting of the oldest blocks. The accumulation of carbonate muds on the distally storm-influenced mid/outer ramp and the differential sedimentation rate of ramp units were strongly influenced by synsedimentary deformations probably related to regional earthquakes. The instability of synsedimentary semiconsolidated carbonate muds on the margin of the slope produced slidings and slumpings of proximal to distal limestone beds passing into debris flows with autoclastic mudstone/limestone breccias triggered by overloading and/or regional tectonic shocks along former active tectonic faults. Our results are consistent with the carbonate ‘debrites’ of the pre-glacial Abenab Formation in Namibia described by Eyles and Januszczak (2007). These authors suggested that these sediments were formed at the foot of steep slopes associated with slope failure accompanying large earthquakes during extension. These deep-water deposits, including the glacigenic Ghaub Formation, are consistent with the regional synrift submarine fan depositional setting proposed by Porada and Wittig (1983), Martin et al. (1985) and Frimmel et al. (2002). Lateral facies and thickness variations show that the series in the northern part of the Sangha aulacogen recording fluvial–lacustrine pre-glacial facies in the Bouenza Formation in RC and Gabon are related to extensional regimes in Central Africa (Alvarez, 1995), rather than open-marine storm-influenced mid/outer ramp facies in the Haut-Shiloango Subgroup in DRC. Alvarez (1995) proposed that a platform or ramp model had developed on a gently sloping shelf located further south. In the southern Lower-Congo aulacogen, Schermerhorn (1974) described the Alto-Shiloango Subgroup in Angola as consisting of pre-flysch turbidites passing to shallow-water sediments, and terminating in carbonates. These carbonates consist of 25–140 m-thick alternations of grey-green to grey-blue limestones and grey to black limestones (Stanton et al., 1963), which accord with the lithological and sedimentary descriptions of the Haut-Shiloango carbonates. The stratigraphic relationships between these carbonates and the overlying glacigenic sediments and melting ice cap series are well established in RC (Alvarez and Maurin, 1991; Alvarez, 1995) and DRC (Cahen, 1950, 1954, 1978; Lepersonne, 1974; Tack et al., 2001; Frimmel et al., 2006). However, no sedimentary evidence of glaciogenicity pointing to deposition during an interglacial–glacial–interglacial period has been clearly identified. Moreover, Schermerhorn (1974) suggested that the Neoproterozoic Upper Tilloids from northern Angola and DRC were deposited as mudflows by slumping of unconsolidated sediments on a shelf, reflecting an abrupt tectonic basinal subsidence. Our new observations carried out during a field reconnaissance in the dry 2011 season confirm that the uppermost Haut-Shiloango Subgroup, the Upper Diamictite Formation and the lower Schisto-Calcaire Subgroup are indicative of deep-water deposits including respectively distal outer ramp carbonate facies, debris-flows, storm-influenced cap carbonates and calciturbidites. According to Eyles and Januszczak (2004), the glaciations may have been influenced by the tectonic uplift of mountains, rift shoulders in rifted basins and newly formed passive margins. The intimate relationship between glaciations and tectonic activity has been recently demonstrated by glacially influ-

enced marine strata dominated by mass flow diamictites and turbidites in a syn-rift system (Strand, 2012). However, Kröner and Correia (1973) interpreted the Upper Diamictite Formation as glaciomarine deposits and the uppermost Haut-Shiloango Subgroup as ‘‘formed in a similar manner as the breccia underlying the continental Dwyka tillite [. . .] considered to be a result of glacial scouring’’. This interpretation is inconsistent with our data. Moreover, the occurrence of glacial deposits is classically associated to sedimentary features such as subglacially striated pavements, facetted and striated clasts, dropstones and far travelled extrabasinal clast assemblages (Boulton, 1978; Etienne et al., 2007; Arnaud and Etienne, 2011). Here, the lack of glacial features such as striated and ‘overdeepened’ sub-basement, glacially shaped clasts akin to bullets, and common extrabasinal fragments within the Upper Diamictite Formation in DRC is inconsistent with diamictites deposited immediately under ice sheets. Nonetheless, we cannot rule out that the Sh8 sediments were pre-Marinoan glacially-influenced by elevated latitude regional glaciers which formed after the rift–drift transition on the passive continental margin.

8. Conclusions Our re-examination of the uppermost Haut-Shiloango Subgroup in DRC leads us to conclude that the Sh8h and Sh8i members do not consist of stromatolites or stromatolitic bioherms as previous studies have suggested. In the present study, based on recent field observations, we have defined new lithofacies, and carried out complementary geochemical analyses, which have led us to reinterpret the depositional environments that occurred between the Sturtian and Marinoan glacial episodes. The lithofacies of Sh8 and Sh8i members display: (i) open-marine synsedimentary features (hummocky cross-laminations and distal tempestites); (ii) synsedimentary contorted structures, sliding and slumping of semi-consolidated limestone beds; (iii) synsedimentary faults (intraformational truncation surfaces); and (iv) autoclastic mudstone breccias of debris flow origin. The vertical stacking of the lithofacies records a deepening-upward evolution from storm-influenced sediments in mid/outer-ramp environments to quiet deep-water settings in distally steepened or sloped ramp environments. The trigger mechanism of the synsedimentary deformations (autoclastic breccias, slumps, etc.) is attributed to tectonic events with active differential tilting and uplifting of blocks in a graben-like basin, related to the 750–670 Ma oceanic spreading in the central-southern Macaúbas Basin, and pre-collisional magmatic arc in the Araçuai-West Congo orogen. Uniform flat non-marine shale-normalized REE+Y patterns record freshwater influences in the Sh8h carbonates. Moderate Y, Zr and Rb contents reflect continental detrital inputs in nearshore environments rather than in deep-water environments. The nearshore sediments were reworked from ’shallow’ inner- to mid-ramp settings in deep-water slope and outer-ramp environments, during the rift-drift transition in the basin. The pre-Marinoan deep-water deposits of Sh8h and Sh8i members are consistent with carbonate ‘debrites’ of the pre-glacial Abenab Formation in Namibia, which were glacially influenced by tectonic uplift of mountains with local formation of elevated latitude glaciers, taking place after the rift–drift transition on a passive continental margin. In our study, on the basis of sedimentary structures, we reinterpret pre-glacial Marinoan deposits as shallow non-marine, or marine freshwater-influenced sedimented reworked from nearshore environments into the deep-water settings along tectonically active continental margins, locally influenced by elevated altitude glaciers. Deeper insights into pre-glacial sedimentary structures associated with Neoproterozoic glacigenic

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