Petrographic and geochemical characterization of the Lower Transvaal Supergroup stromatolitic dolostones (Kanye Basin, Botswana)

Accepted Manuscript Petrographic and geochemical characterization of the Lower Transvaal Supergroup stromatolitic dolostones (Kanye Basin, Botswana) Fulvio Franchi PII: DOI: Reference:

S0301-9268(17)30546-6 https://doi.org/10.1016/j.precamres.2018.02.018 PRECAM 5029

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Precambrian Research

Received Date: Revised Date: Accepted Date:

4 October 2017 19 February 2018 25 February 2018

Please cite this article as: F. Franchi, Petrographic and geochemical characterization of the Lower Transvaal Supergroup stromatolitic dolostones (Kanye Basin, Botswana), Precambrian Research (2018), doi: https://doi.org/ 10.1016/j.precamres.2018.02.018

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Petrographic and geochemical characterization of the Lower Transvaal Supergroup stromatolitic dolostones (Kanye Basin, Botswana) Fulvio Franchi1 1. Department of Earth and Environmental Science, Botswana International University of Science and Technology (BIUST), Private Bag 16, Palapye, BOTSWANA

ABSTRACT The 2.5 Ga stromatolitic dolostones from the Lower Transvaal Supergroup in the Kanye Basin (Botswana) pre-dates the first iron deposits, recording conditions before the great oxidation event (GOE). These dolostones have been deposited within a shallow marine carbonate platform extending from Zimbabwe to South Africa. The Lower Transvaal Supergroup carbonates of the Kanye Basin have been affected by the circulation of metasomatic fluids related to emplacement of the Moshaneng Dolerites (1.9-2.1 Ga). Here, geochemical and petrographic characterization of the Lower Transvaal Supergroup stromatolites is presented to shed light onto i) the effect of metasomatic fluids on the geochemistry of ancient carbonates, and ii) the environmental conditions prevailing before the onset of the GOE in the epeiric seas along the western margins of the Kaapvaal Craton. The dolomites show high Fe an Mn contents (average 2000 ppm and 3500 ppm, respectively) and very low Na contents. The overall rare earth elements (REE) pattern of dolomite vary consistently across the different dolomite facies with a sensible increase of ∑REE in the altered dolomites. The overall REE patter lacks La, Ce and Gd anomalies (average 1, 0.91, 1.05 respectively) and shows an overall chondritic Y/Ho ratio. The Eu anomaly is slightly negative or absent in most of unaltered samples. A positive Eu anomaly (average 1.28) and an overall ∑REE enrichment have been detected in samples altered by metasomatic fluids.

The dolomite samples from Ramonnedi Formation show Fe and Mn enrichment typical of Precambrian carbonates. Petrographic and geochemical analyses reveal that dolomitization and incipient diagenesis have probably affected the platform carbonates inducing neomorphism and recrystallization (i.e. xenotopic and syntaxial dolomite). The dolomitization processes have mobilized Y giving rise to a near chondritic Y/Ho ratio. Evidences for circulation of hydrothermal fluids related to the intrusion of Moshaneng Igneous Complex and Moshaneng Dolerites are shown only in one sample. Most samples of stromatolitic dolostone have negative Eu anomaly suggesting deposition in a closed basin system restricted from open marine circulation.

Key words: Archean carbonates, Transvaal Supergroup, carbonate platform, stromatolite, dolomite, rare earth elements

1. INTRODUCTION The geochemical composition of carbonates has been often used as proxy for the investigation of geochemical signatures of early life in Precambrian oceans. In the last two decades new analytical approaches in carbonate petrology have improved understanding of the interaction between biological activity and carbonate precipitation (e.g., Webb and Kamber, 2000; Kamber and Webb, 2001; Van Kranendonk et al., 2003; Bolhar et al., 2004; Kamber et al., 2004; Nothdurft, et al., 2004; Bau and Alexander 2006; Bolhar and Van Kranendonk, 2007; Shields, 2007; Allwood et al., 2010; Riding et al., 2014; Franchi et al., 2015; 2016; Petrash et al., 2016). Particularly the shale-normalized (SN) pattern of rare earth elements (REE) and Y, and the distribution of redox sensitive trace elements from ancient microbial carbonates can provide crucial paleoenvironmental information and have been used to investigate ancient seawater chemistry (e.g., Holocene, Webb and Kamber, 2000; Devonian, Nothdurft et al., 2004; Franchi et al., 2015; 2016; Archaean, Kamber and

Webb, 2001; Petrash et al., 2016). Nevertheless, the evaluation of the effects of diagenesis on the pristine composition of ancient carbonates has proven to be challenging (e.g., Brasier et al., 2002; Schopf et al., 2002; Van Kranendonk et al., 2003). Several studies have demonstrated that diagenetic mobilization of trace elements, neomorphism and contamination from exogenous sources can alter the original geochemical signature of marine water (e.g., Bau and Dulski, 1996; Reynard et al., 1999; Shields and Stille, 2001; Johannesson et al., 2006; Petrash et al., 2016). Precambrian carbonates are often silicified and recrystallized (Fedo et al., 2001; Allwood et al., 2006; Kamber and Webb, 2007) to the point where microstructures are obliterated and therefore can lead to misinterpretation of mineralization pathways. In these ancient carbonates, isotopic signatures of carbon and oxygen are blurred and do not provide unambiguous paleoenvironmental proxies (e.g., Tang et al., 2016 and references therein). On the other hand, REE are considered to be extremely stable during diagenesis and recrystallization of carbonates as they substitute Ca2+ within the lattice (e.g., Banner et al., 1988; Zhong and Mucci, 1995; Kamber and Webb, 2001; Bau and Alexander, 2006) and may provide an important geochemical tool for the reconstruction of ancient environments. A broad consensus developed over the last few decades suggests that, although original micro-fabrics might be obliterated during diagenesis, the effect of diagenetic fluids circulation does not obliterate the geochemistry of the carbonates (e.g., Van Kranendonk et al., 2003; Franchi et al., 2016 and references therein), which can be still used as proxies to aquatic environments (e.g., Kamber and Webb, 2001). Is this always true? This question underpins the current research aimed at discriminating between the effect of diagenesis, and particularly circulation of metasomatic fluids, and the original composition of platform carbonates precipitated before the Great Oxidation Event (GOE). Assuming that the final goal of the research is to assess the environmental condition at the time of deposition, including the effect of O2 on the distribution of trace elements and REE, it is paramount to assess the effect of diagenetic fluids on the geochemistry of the Archean dolomites. It is likely that post depositional processes, such as circulation of metasomatic

fluids as well as dolomitization itself, have affected the geochemical signature of microbial carbonates to the extent that original patterns have been overprinted (Webb and Kamber, 2000; Nothdurft et al., 2004; Franchi et al., 2015; 2016). The analysis of a carbonate platform, restricted and constrained spatially and temporally, may provide a perfect laboratory for testing the effects of depositional and post depositional processes on the geochemistry of the carbonates. A carbonate platform, in fact, records the complex interactions between biosphere, ocean chemistry and geological processes such as sea level fluctuations and basin subsidence (e.g., Bolhar et al., 2005; Sumner and Beukes, 2006). Detailed studies of platform and shelf carbonates are available for Neoarchean successions of the Campbellrand/Malmani subgroups of the Transvaal Supergroup in South Africa (Kamber and Webb, 2001; Knoll and Beukes, 2009 and references therein; a focus on the Gamohaan and Frisco formations is presented in Sumner, 1997; Eroglu et al., 2015), the Early Paleoproterozoic Turee Creek Group of Western Australia (Martindale et al., 2015 and references therein) and the Fennoscandian Shield (McLoughlin et al., 2013 and references therein). The samples under investigation in this work come from the Taupone Dolomite Group in the Lower Transvaal Supergroup of Botswana (Kanye Basin; Figs. 1-3) considered to be correlative to the Campbellrand/Malmani carbonates in South Africa. To date relatively few studies have focused on the Lower Transvaal carbonates of the Kanye Basin, despite their potential importance for understanding Archean and Paleoproterozoic seawater evolution and environmental change across the GOE. Archean and Proterozoic carbonate platforms are often associated with iron formation and might be interesting for i) their economic potential and ii) for the paleoenvironmental reconstruction of the marine basin before and during the GOE. The geochemical composition of these sediments (including redox sensitive trace elements and REE) may preserve signature of ancient life developed under extreme chemo-physical conditions. The study of Precambrian marine sediments might therefore shed light on the conditions that

underpinned the development of early life as well as the effects that these microbial communities exerted on the ancient atmosphere and hydrosphere. Particularly, the geochemical characterization of marine sediments deposed at the Archean/Paleoproterozoic transition may contribute to better understanding the processes that led to the GOE (e.g., Johnson et al., 2013). During the GOE the Earth experienced a rise in atmospheric O2 developing a stable surface ocean oxic conditions for the first time about 2.45-2.33 Ga (e.g., Farquhar et al., 2000; Bekker et al., 2004; Luo et al., 2016). Although it has been proposed that the metabolism of oxygen-producing photoautotrophs led directly to the GOE’s rise in atmospheric oxygen (Kopp et al., 2005; Knoll and Beukes, 2009; Johnson et al., 2013), the role of phototrophs before the GOE is still unclear (Des Marais, 2000). The present study reports geological, petrographic and geochemical (trace elements including REE+Y) data from stromatolitic dolostone from the Ramonnedi Formation (Taupone Dolomite Group) in order to answer the following questions: i) how are pristine signatures of the environment preserved in ancient carbonates that may record different degrees of oxygenation, and what was the role of photoautotrophs before the GOE? ii) To what extent diagenesis and metasomatism can affect the distribution of trace elements and REE in Precambrian carbonate sediments?

2. GEOLOGICAL BACKGROUND The Transvaal Supergroup occurs in three basins lying unconformably on the crystalline rocks of the Kaapvaal Craton (Fig. 1): the Transvaal Basin and the Griqualand West Basin, in South Africa, and the Kanye Basin, in southeastern Botswana (Eriksson et al., 2001; 2006; Bumby et al., 2012). In the Kanye Basin the Transvaal Supergroup rocks consist of a thick sedimentary sequence that was deposited during a period from 2.7 to 2.1 Ga (Catuneanu and Eriksson, 1999). The early Proterozoic Transvaal Supergroup in Botswana unconformably overlies the Kanye Volcanic Formation and the Gaborone Granite (Bumby et

al., 2012), and the Lobatse Volcanic Group (Mapeo and Wingate, 2009) (Figs. 3-4). The Transvaal Supergroup sequence predates the ca. 2.05 Ga Bushveld Complex igneous event within the Kaapvaal Craton (e.g., Eriksson et al., 1995; Buick et al., 2001; Bumby et al., 2012). In the Kanye Basin the Moshaneng Igneous Complex (ca. 2.10 Ga, Aldiss et al., 1989; 2.05 Ga, Mapeo et al., 2004) is considered to be the equivalent of the Bushveld Complex of South Africa (e.g., Mapeo et al., 2006). The Moshaneng Igneous Complex is made of mainly granioids rocks occurring around Moshaneng village (Fig. 2A-B). Three main units have been described: i) porphyritic granites; ii) diorites with associated mafic rocks; iii) syenite (Aldiss et al., 1989; Mapeo and Wingate, 2009). Generally, the complex shows a concentric pattern with a dioritic core surrounded by granitic rocks. The porphyritic granites and diorites of the Moshaneng Igneous Complex have intruded the Lower Transvaal dolomites inducing localized hydrothermal alteration and metasomatism of the carbonates and the consequent formation of talc (Aldiss et al., 1989). The Moshaneng Dolerites have then intruded both the Transvaal dolomites and the Moshaneng Igneous Complex (Mapeo and Wingate, 2009), around 1927 Ma (Hanson et al., 2004), inducing a second stage of metasomatic alteration.

2.1. Lower Transvaal Supergroup The ca. 2.64-2.43 Ga Lower Transvaal Supergroup (Mapeo et al., 2006), which forms the subject of the present study, consists of a mixed siliciclastic-carbonate ramp that grades upward into an extensive carbonate platform, overlain by subtidal banded iron-formation (e.g., Sumner and Beukes, 2006). The Lower Transvaal Supergroup in the Kanye Basin is represented by the siliciclastic sediments of Black Reef Quartzite Formation and by the carbonate platform of the Ramonnedi Formation covered by the Masoke and Kwakgwe Chert Breccia formations (Mapeo et al., 2006; Bumby et al., 2012).

The Black Reef Quartzite Formation (ca. 2.64 Ga; Mapeo et al., 2006) rests unconformably on the Kanye Felsites and Lobatse Volcanic Group (Fig. 3). The Black Reef Quartzite Formation sediments constitute a fining upward sequence reflecting a progressive increase of the water depth from continental to shallow marine (supratidal to tidal) conditions. Aldiss et al. (1989) suggested deposition within tidal channels along a clastic coastline where the fine-grained sediments have been winnowed away from gravels and sands. The transition between the Black Reef Quartzite Formation and the overlying Taupone Dolomite Group carbonate platform is gradational (Fig. 4). This transition and the general deepening upward trend are ascribable to a first order transgressive cycle (Aldiss et al., 1989). The Taupone Dolomite Group is considered correlative to the Chuniespoort Group in South Africa, and the Ramonnedi Formation is considered to be correlative to the Campbellrand and Malmani subgroups in South Africa (e.g., Mapeo et al., 2006). These units represent a carbonate platform that covered the whole extension of the Kaapvaal Craton, including the Kanye Basin (Fig. 1) (Beukes, 1987; Knoll and Beukes, 2009). In South Africa the carbonate platform sequence is considered to be younger than the 2.6-2.7 Ga Vetersdorp lava (U/Pb zircon age; see Beukes, 1987) and older than the 2.2 Ga Pretoria Group lava (Rb/Sr age; Hamilton, 1977). Tuff beds within mainly carbonate-facies units in the Campbellrand Subgroup (Griqualand West basin) have yielded dates ranging between 2.59 and 2.55 Ga (Barton et al., 1994; Altermann and Nelson, 1998). Overlying the carbonates of the Campbellrand Subgroup in the Griqualand West Basin are the 2.49 and 2.43 Ga (Nelson et al., 1999; Pickard, 2003) shales and banded iron-formation of the Kuruman and Griquatown formations. Interestingly, Eroglu et al. (2015) consider the Campbellrand-Malmani platform to have been stable before and during the GOE (2.4-2.32 Ga). The stratigraphy and sedimentology of the Campbellrand-Malmani platform has been the subject of numerous studies (e.g., Beukes, 1977; 1987; Sumner and Grotzinger, 2004; Sumner and Beukes, 2006; Knoll and Beukes, 2009; Schröder et al., 2009; Eroglu et al.,

2015). The platform is characterized by two main lithofacies: a shallow water stromatolitic sequence and a basinal non-stromatolitic sequence of laminated to massive carbonates, shale and jaspilitic chert (Knoll and Beukes, 2009 and references therein). The original limestone has been entirely replaced by diagenetic dolomite (Beukes, 1987) and the carbonates, at least on the South Africa side of the platform, have undergone lower greenschist facies metamorphism (e.g., Miyano and Beukes, 1984; Knoll and Beukes, 2009). The Campbellrand-Malmani carbonate platform extends into the Kanye Basin and is formalized as the Taupone Dolomite Group (Fig. 3). The lowermost part of the Taupone Dolomite Group is the Ramonnedi Formation (Figs. 3-4). This formation consists of shales interbedded with quartz arenites and chert-poor dolomite (Fig. 4). The Ramonnedi Formation commences with bluish-grey shales interbedded with dolomite and passes to brownish dolomites interbedded with bluish-grey shales and quartz arenite (Figs. 5A-F). The whole succession is capped by bluish-grey well bedded dolostone. Toward the top of the formation dolerite dikes become frequent and the dominant lithotype is chert-rich dolomite (Fig. 5G) and locally oolitic chert (Fig. 5H). Toward the top of the Ramonnedi Formation the dolomite is associated with chert replacements, this change in lithotype is formalized as the Moshaneng Member (Aldiss et al., 1989; Fig. 3). This member consists of ‘ribbon-like’ chert interlayered with bluish-grey dolomite (Fig. 5I). The uppermost units of the Taupone Dolomite Group are the ironstones of the Masoke Formation and the jaspilitic chert of the Kgwakgwe Chert Breccia Formation (Carney et al., 1994; Mapeo, 1997) (Figs. 3-4). The Masoke Formation is divided into two main lithotypes: i) reddish layered chert with dolomite interbeds; ii) banded ferruginous chert (banded iron formation; BIF). The Kgwakgwe Chert Breccia unit is made up by reddish-brown chert breccia dominated by jaspilitic fragments cemented by reddish-brown jaspilitic material. The Masoke Formation, together with the Kgwakgwe Chert Breccia, is considered to be an equivalent of the BIF (Kuruman Iron Formation and lower Penge Iron Formation; Sumner

and Beukes, 2006) that marks the top of the carbonate platform sequence in South Africa (Carney et al., 1994; Eriksson et al., 2006; Knoll and Beukes, 2009; Bumby et al., 2012).

3. METHODS The Ramonnedi Formation carbonates are unevenly exposed in the study area (Fig. 2). Field work has been carried out in summer 2015, 2016 and winter 2017. Sampling has been carried out along a composite NE-SW section from Tletlesi and Mokalaka ridges, north from Moshaneng village, to the Diphawana hills south of Kanye village (Fig. 2). For sake of clarity here we use the term dolomite with reference to the mineral and the term dolostone with reference to the carbonate rock dominated by the mineral dolomite. To distinguish between the lithofacies close to the intrusion, allegedly skarn deposits, and the stromatolitic dolostone the descriptive terms ‘altered’ and ‘unaltered’ facies will be used throughout. Dolomite fabrics have been described following the classification by Sibley and Gregg (1987) (see also Warren, 2000). Thirty-three (33) samples of dolostone, sandstone and chert were collected and prepared for geochemical and petrological analysis. Standard thin sections for petrographic analysis have been prepared at the University of Botswana (UB) and the Council of Geoscience (South Africa). Eighteen (18) samples of carbonates have been selected for further petrographic and geochemical analyses. Thin section blanks were sub-sampled for geochemical analyses (Supplementary Materials 1 and 2) using a hand-held drill following the method reported by Martindale et al. (2015). Morphological analyses have been carried with a JEOL JSM-7100F field emission scanning electron microscope (FE-SEM) equipped with an electron dispersive spectrometer (EDS) at Botswana International University of Science and Technology (BIUST).

X-ray diffraction (XRD) analyses were performed with a Bruker D8 Advance X-ray diffractometer (Cu Kα X-ray source) at BIUST. Results are reported in Table 1. Major elements were analyzed by X-ray fluorescence (XRF) at the Botswana Geoscience Institute (BGI) using a PANanalytical Zetium Sequential Wavelength Dispersive XRF with Rhodium anode and equipped with two detectors: low gas detector (methane/argon) and scintillation detector. 1.2 g of powders were weighed on an analytical balance and mixed with 7.2 g of a fused, anhydrous mixture constituting of 49.75% Li2B4O7, 49.75% LiBO2 and 0.50% LiBr. The mixture was then heated until molten in a Phoenix Automated Fusion machine VFD6000 and ultimately became glassy fused beads. All elements were measured using K-lines (Kα1, 2). The conversion of measured X-ray intensities into element concentrations was based on calibration of the XRF spectrometer by measuring the peak intensities for a series of reference materials (12 CRMs). Results are reported in Table 2. Trace elements including REE (Table 2) were analyzed with inductively coupled plasma mass spectrometry (ICP-MS) at Activation Laboratories (Canada) following the protocol proposed by Rongemaille et al. (2011). Ca. 5 mg of powdered sample were leached with 5% v./v. acetic acid for 24 h and then analyzed with a Perkin Elmer Sciex ELAN 6000 ICP-MS. The relative abundances of REE were normalized to the Post Archean Australian Shale (PAAS; Taylor and McLennan, 1985), representing the estimated average terrigenous input to the oceanic environment. ΣREE, Y/Ho, La, Ce, Eu and Gd anomalies (Table 3) and the relation between these proxies and trace and transition elements were calculated for all the studied samples. Light REE (LREE) fractionation was calculated as PrSN/YbSN to avoid bias due to anomalous La and Ce concentrations. Medium REE (MREE) enrichment with respect to heavy REE (HREE) was calculated as GdSN/YbSN. Normalized La, Ce, Eu and Gd anomalies were calculated using the geometric equation given by Lawrence et al. (2006): (La/La*)SN = LaSN / [(3 * PrSN) - (2 * NdSN)]

(1)

(Ce/Ce*)SN = CeSN / [PrSN * (PrSN / NdSN)]

(2)

(Eu/Eu*)SN = [EuSN / [(SmSN)2 * TbSN]1/3

(3)

(Gd/Gd*)SN = [GdSN/ [(TbSN)2 * SmSN]1/3

(4)

4. RESULTS 4.1. Lithofacies and petrographic characterization Based on field observation five dolostone lithofacies can be distinguished: massive dolostone, stratiform stromatolite, columnar and micro-columnar stromatolite and domal stromatolite (Fig. 5B-F). The bottom of the carbonate sequence is characterized by massive dolostone (FK_5B) interbedded with quartz arenite and siltstones (Fig. 5A) showing clear evidence of desiccation cracks. This mixed carbonate siliciclastic unit grades upward to domal stromatolite and stratiform stromatolite with wrinkled laminae (FK_8A-B) (Fig. 5B-C). Stratiform stromatolites are characterized by flat or wrinkly lamination showing an alternation of sub-millimeter dark laminae and slightly thicker whitish to light grey laminae (Fig. 5C-D). The size of the stromatolite increases upward where large domal stromatolites become dominant (FK_8C, FK_8E) (Fig. 5E). The large domal stromatolites (stromatolite mounds) grade laterally into stratiform stromatolite (sensu Coffey et al., 2013) overlaid by microcolumnar stromatolites (FK_8D, FK_30) (Fig. 5C-D) and columnar stromatolites (FK_8F) (Fig. 5F). The Ramonnedi Formation sequence of stromatolitic dolostone ends abruptly with the first occurrence of laminated chert (Fig. 5G). The lowermost part of this chert rich member is characterized by several levels of oolitic chert (FK_18, FK_27) (Fig. 5G-H). Locally, dolomite has been replaced by quartz, especially along macroscopic lineations produced by the original stromatolite lamination (Fig. 5I). The mineralogical composition of the stromatolitic dolostone is mainly dolomite and quartz (Table 1). The diffractograms of dolomite-dominated samples exhibit a variable carbonate

phases, with cell parameters and d-spacing (104) suggesting the presence of two end members, namely calcian dolomite and ankerite. The samples collected close to the intrusion have a higher quartz/dolomite ratio and are characterized by the presence of calcite, dolomite, talc, diopside, chrysotile and probably forsterite and tremolite (Table 1). Close to the intrusions the carbonates are almost entirely replaced by talc and other secondary minerals locally cross cut by veins of chrysotile (Table 1). Six different carbonate microfacies (Fig. 6) can be identified from the unaltered facies: i) inequigranular mosaic of xenotopic turbid dolomite (anhedral crystals) (Fig. 6A); ii) xenotopic turbid dolomite with quartz pockets; iii) inequant mosaic of hypidiotopic turbid dolomite with microcrystalline cements (Fig. 6B); iv) idiotopic rhombohedral dolomite (Fig. 6C); v) microcrystalline dolomite (Fig. 6D); vi) coarse neomorphic dolomite with syntaxial replacements (Fig. 6F). Xenotopic and hypidiotopic dolomite is normally medium to coarsely crystalline (Folk, 1962). Dolomite types with uneven crystal boundaries are by far the most abundant fabric (i.e., xenotopic dolomite). Locally the inequigranular mosaic of dolomite is characterized by microcrystalline dolomite, microcrystalline quartz and patches of replacive quartz (Fig. 6D). The microcrystalline dolomite shows evidence of recrystallization to rhombohedral dolomite (Fig. 6D). The rhombohedra of dolomite are often characterized by turbid centers with clear margins (Fig. 6B-D). The euhedral crystals form either a planar mosaic of idiotopic dolomite (Fig. 6B) or an inequigranular planar mosaic with a microcrystalline dolomite matrix (Fig. 6C-D). Quartz occurs either in patches of as replacive quartz (Fig. 6D) or in thin layers of brownish, turbid microcrystalline quartz (Fig. 6E). In the oolitic lithofacies (Fig. 5H) the ooids have carbonate cortices and are cemented by microcrystalline to coarse quartz (Fig. 6G-H). The ooids consist of radiaxial crystals of dolomite partially replaced by quartz (Fig. 6H).

4.2. Geochemical composition

4.2.1. Major and trace elements The dolostone samples from the Ramonnedi Formation show different rates of silicification (SiO2 between 1.16% and 98.65%; Table 2). Al2O3 content varies from 0.03% and 4.17% showing no direct correlation with SiO2 (Table 2). CaO and MgO content varies among the different samples (1.03-39.17% and 1.04-30.05%, respectively). XRF analyses revealed an overall low content of other major oxides such as K2O, TiO2, MnO and Fe2O3 (Table 2). The average Ca/Mg ratio (ppm/ppm) of the unaltered dolomite is ca. 2, while it is as high as 9.35 in the altered facies (Table 3). The average Mg content in the altered facies is as low as 26000 ppm, less than half of the Mg content observed in the unaltered dolomite. The average Ca content shows a threefold increase from unaltered to altered dolomite (Table 3). ICP-MS analyses reveal a general enrichment of Mn in all samples with an average of 3267 ppm and 4728 ppm from unaltered and altered facies, respectively. The average V content decreased from 1.6 ppm in the unaltered dolostone to 0.3 ppm in the altered facies. A similar trend is shown by Fe, Mo and Ni (from 2593 ppm to 143 ppm, from 13 ppm to 0.1 ppm and from 467 ppm to 36 ppm, respectively). Samples from altered facies are generally enriched in As (average 2 ppm), Se (average 2 ppm), Sr (average 126 ppm), I (average 4 ppm) and Ba (average 15 ppm). These elements are generally one order of magnitude less concentrated in the unaltered dolostone. Lead is slightly enriched in the altered facies with average of 10.7 ppm, in contrast to the 7.2 ppm measured in unaltered dolostone. Concentrations of elements typically associated with siliciclastic material are very low (Table 2): Rb (average 0.7 ppm), Cs (average 0.1 ppm), Zr (average 0.2 ppm), Hf (average 0.0 ppm), Th (average 0.2 ppm).

4.2.2 REE distribution

All but one sample (FK_23) show near-chondritic Y/Ho ratios (average 32; Table 3). The massive dolostone of sample FK_23 shows super-chondritic Y/Ho (76.32) (Fig. 7). The REE pattern varies from unaltered dolostone to altered facies (Fig. 7). The average ∑REE content from the unaltered samples is 7.6 ppm (Table 3). The normalized REE distributions from unaltered dolomite samples are quite consistent among the different lithofacies (namely massive dolomite, stratiform stromatolite, domal stromatolite and columnar stromatolite) showing a general slight LREE depletion with respect to HREE (average PrSN/YbSN = 0.91) and a slight MREE enrichment (average GdSN/YbSN = 1.34) (Fig. 7). Only massive dolostone (FK_03 and FK_32) and mudrocks (FK_05B) show a bulged pattern with strong LREE depletion (Figs. 7A, 8). The samples of unaltered dolomite show slightly negative La, Ce and Gd anomalies (0.93, 0.89 and 0.91, respectively) whereas the Gd anomaly is slightly positive (1.06) (Table 3). Samples from the altered facies show an average ∑REE content of 20 ppm and an overall flat REE pattern (Fig. 7). Only sample FK_33 shows a sensible enrichment in LREE (light REE) with respect to HREE (Table 3; Fig. 7). This sample shows a positive anomaly of La and Eu (average of 1.69 and 3.33, respectively) while no Ce is present (Table 3). Other samples from altered facies lack a real La anomaly and have a slightly negative Eu anomaly (average 0.88), whereas Ce and Gd anomalies were not found (average of 0.95 and 1.02, respectively). The five sub-sample collected from the oolitic chert (FK_18) show an overall flat REE pattern (PrSN/YbSN = 0.97) with a slight enrichment of MREE compared to HREE (GdSN/YbSN = 1.21) (Table 3; Fig. 7F). This lithofacies shows slightly negative Ce, Eu and Gd anomalies (0.90, 0.92 and 0.93, respectively). Because of the weak acid leaching procedure these values are considered to be representative of the dolomitic cortices of the oolites.

5. DISCUSSION

5.1. Evolution of the Lower Transvaal Supergroup carbonate platform The 2.6-2.4 Ga old carbonates of the Ramonnedi Formation conformably overlie the earliest Neoarchean deposits (i.e. BRQ, Lobatse Volcanic Group; Fig. 2). The Ramonnedi Formation predates the chert that records the onset of the GOE in the Kanye Basin (Mapeo et al., 2006) and probably pre-dates the ca. 2.2 Ga onset of the Lomagundi C-isotopes excursion (e.g., Martin et al., 2013). The carbonates of the Ramonnedi Formation, and stratigraphically equivalent units from the Transvaal and Griqualand West basins, have been deposited along a continental shelf that covered most of the Kaapvaal Craton margin (Bumby et al., 2012 and references therein). The Lower Transvaal Supergroup carbonates represent one of the first examples of a carbonate platform that formed along a Neoarchean continental margin. The Ghaap/Chuniespoort groups in South Africa and the Taupone Dolomite Group in Botswana play a crucial role in the reconstruction of the pre-GOE conditions across the late Neoarchean to early Paleoproterozoic transition (2.65–2.40 Ga; e.g., Beukes and Gutzmer, 2008), because they provide a continuous stratigraphic record of a well-preserved carbonate platform capped by iron formations (e.g., Knoll and Beukes, 2009). The bottom of the Ramonnedi Formation shows stacked layers of mudrock, cross laminated siltstone and quartz-arenite (locally arkose), and massive dolostone (Figs. 2, 5). The abundance of mudrock interbedded with planar symmetrical ripple cross-laminated sandstone beds and siltstone with evidence of desiccation cracks most likely represents deposition along a tidal flat. The bulk of the Ramonnedi Formation is characterized by a thick sequence of stromatolitic dolostone with lithofacies that vary from domal and stratiform at the base to large domal and columnar at the top of the sequence (Fig. 4). The large domal forms (Fig. 5E) are analogues to the giant stromatolites of the Campbellrand/Malmani platform (Knoll and Beukes, 2009 and references therein) and are similar to the unusual meter-scale stromatolitic buildups described in the Kazput carbonates (Western Australia; Martindale et al., 2015). These larger stromatolite domes resulted from protracted growth between episodes of sand deposition (Martindale et al., 2015). The wrinkled laminae typical

of stratiform stromatolites (Fig. 5C-D) have been interpreted elsewhere as evidence of matbuilding microbial communities (e.g., Martindale et al., 2015). The Ramonnedi Formation carbonate platform grades up into a sequence of chert and jaspilitic chert (Figs. 3-4). The base of this unit is marked by the occurrence of oolitic chert (Fig. 5H). Ooids associated with stromatolites are often considered to be evidence for peritidal to subtidal conditions (Coffey et al., 2013; Martindale et al., 2015; Petrash et al., 2016) and have been described by Sumner and Grotzinger (2004) from the CampbellrandMalmani carbonates (cf. Fig. 8 in Sumner and Grotzinger, 2004). Chert precipitation has probably occurred in the littoral mixing zone (e.g., Knauth, 1979, Back et al., 1986). The topmost sediments of the Group are considered to be correlative to the BIF described at the top of Chuniespoort Group (i.e., Penge Iron Formation) in South Africa and believed to be deposited after a major transgression (e.g., Bumby et al., 2012). In conclusion, the Taupone Dolomite Group sequence begins with the BRQ interpreted as tidal channel fill sequence (Aldiss et al., 1989), grades into a well-developed carbonate platform and is capped by oolitic units and by banded chert probably deposited under subtidal conditions. The bottom of the Ramonnedi Formation therefore represents the first stage of transgression with transition from a shallow marine environment (between fair wave base and high tide level) to a low lying tidal mud flat and to a protected lagoon characterized by evaporative processes, while the chert (and subsequent BIF of the Masoke Formation) marks the drowning of the carbonate platform.

5.2. Petrography and genesis of the Ramonnedi Formation dolostones The Ramonnedi Formation dolomite clearly represents a diagenetic product as shown by the crystal fabrics, dominated by non-planar xenotopic and hypidiotopic dolomite. Microcrystalline dolomite is scanty in the studied samples and is mostly preserved as intercrystalline phase in coarser dolomite planar fabrics (Fig. 6C). The paucity of pristine

microcrystalline dolomite, together with the dominance of coarse grained dolomite, is indicative of recrystallization (dolomitization) of the original platform carbonates. The dolomite samples, especially the microfacies dominated by xenotopic and hypidiotopic dolomite mosaics (Fig. 6A-B), show high inter-crystalline secondary porosity. The replacement of CaCO3 through dolomitization leads to an overall decrease of the volume of the rock (about 12%) accompanied by an increase in porosity (Van Tuyl, 1918; Warren, 2000 and references therein). To date no unambiguous genetic model for the Ramonnedi Formation exists. Eriksson and Warren (1983) considered the Malmani dolomite in South Africa to be an example of ancient dolomitization that occurred during meteoric-seawater mixing. This model implies that mixing of two waters that are saturated with a carbonate phase will lead to a solution that may be undersaturated with respect to that particular phase and saturated with respect to another phase, dolomite in this case (Warren, 2000). When CO2-saturated meteoric groundwaters mix with normal seawater the mixture will maintain the same Mg/Ca ratio but, because of the lowered concentrations of sulfates, and the increased proportion of carbonates over bicarbonates in the solution, dolomite is more likely to precipitate (Lippmann, 1973; Folk and Land, 1975; Warren, 2000). Crystals of dolomite precipitated under meteoric mixing zone conditions may show a turbid center with clear margins like the rhombohedra in figure 6C. The turbid centers would be generated by incipient alteration of low-Mg calcite precursor and may preserve micro-inclusions of such minerals and micro-cavities (due to calcite dissolution) even after dolomitization (Warren, 2000). In general, the dolomite rhombohedra zonation reflects the evolution of pore water from (near) saturation to under-saturation in Ca (Sibley, 1980). According to Gregg and Sibley (1984), mosaics of idiotopic dolomite with planar crystal boundaries indicate growth temperatures below 50℃, while xenotopic dolomite may results from higher temperatures (>50℃). Both planar and non-planar

dolomite can form as a cement, by replacement of limestone, or by neomorphic recrystallization of a precursor dolomite (Gregg and Sibley, 1984; Sibley and Gregg, 1987). The circulation of metasomatic fluids through the Ramonnedi Formation sediments have induced the replacement of dolomite by talc, tremolite chrysotile and other alteration minerals (Table 1) localized close to the intrusion. The circulation of fluids may have also affected the Ca/Mg ratio of the carbonates with partial replacement of dolomite, typical of unaltered facies, with late stage calcite (Tables 1 and 3). The presence of forsterite associated with altered samples (i.e. FK_28B and FK_31C) provides additional evidence for the interaction of the studied rocks with hydrothermal fluids (e.g., Herrero et al., 2011). With increasing pCO2 in fluids, and decreasing temperature, forsterite can form through the reaction of talc and magnesite or, in the case under investigation, most likely by reaction with dolomite. Talc probably formed under the influence of hot hydrothermal fluids, and the later cooling and chemical evolution of this fluids led to the formation of forsterite and late calcite (Warren, 2000). The circulation of these high-temperature fluids induced alteration of dolomites to talc and formation of local asbestos. Considering the micro-fabric and the crystal textures it appears that the carbonates of the Ramonnedi Formation were entirely replaced and overgrown. Due to the widespread abundance of coarse dolomite types, the diagenetic overprinting does not seem to be facies dependent. Silicification, on the other hand, affected preferentially the thin lamination within stratified stromatolites further enhancing the layered nature of these rocks. Petrash et al. (2016) showed how during silicification the metastable microcrystalline dolomite is obliterated, leading to micro-quartz growth, while euhedral dolomite is preserved. This process led to the formation of a complex facies where layers of microcrystalline quartz and rhombohedra of dolomite cut across a matrix made up of microcrystalline dolomite (Fig. 6D). The silicification of carbonates takes place only when pore fluids are supersaturated with respect to silica, and undersaturated with respect to carbonate minerals that become dissolved (Hesse, 1989). These conditions probably existed during the drowning of the

platform, up until before deposition of the Masoke Formation’s BIF when the carbonate platform was blanketed by laminated chert. The replacement of carbonates by silica can be fostered by microbial metabolisms that, by lowering the pH, affects carbonate solubility inducing precipitation of silica (e.g., Hesse, 1989). A sensible lowering of pore water pH can also be induced by mixing of marine and continental meteoric waters (Knauth, 1979; Petrash et al., 2016). Both scenarios indicate that the silicification process must have taken place within the shallow burial diagenesis zone where eustatic sea level changes caused the meteoric water - seawater mixing zone to migrate (Simonson and Hassler, 1996; Petrash et al., 2016). The absence of other mineral replacements, together with the occurrence of microcrystalline quartz, suggest that the process is not strictly related to the circulation of metasomatic fluids. An alternative scenario for the formation of the Ramonnedi Formation dolomite might be the direct precipitation in a restricted evaporitic lagoon. The reduced water exchange within a shallow epeiric sea creates particularly high productivity conditions favorable for phototrophs to thrive in an otherwise hostile global marine environment (Van Kranendonk et al., 2003). The presence of a highly evaporative setting along the margin of the Taupone Dolomite Group platform is consistent with early dolomitization in seawater. Dolomite formation within a lagoon probably began with the precipitation of nanometre-scale dolomite aggregates, possibly aided by the presence of extracellular polymeric substances (EPS) secreted by microbial communities and acting as catalyzer (e.g., Preto et al., 2014). During a later diagenetic stage, planar and non-planar dolomites probably replaced the primary dolomite phase as documented in the Dolomia Principale by Frisia and Wenk (1993) and Frisia (1994). A purely abiotic origin seems therefore unlikely for this dolomite, especially considering the strict association with stromatolites and the presence of putative microbial fabrics (i.e. spherulite, Fig. 6I). Spherulite, or spheroidal dolomite, is commonly considered as primary evidence of bacterial activity in ancient carbonates (Mastandrea et al., 2006; McKenzie and Vasconcelos, 2009).

Whatever genetic scenario is deemed suitable for the formation of the Ramonnedi Formation dolomite, it has to consider that the correlative carbonates of the Campbellrand/Malmani platform in South Africa are only partly dolomitized and mostly preserve calcite, and often aragonite, precursors (Webb and Kamber, 2000; Sumner and Grotzinger, 2004). Hence, the extensive dolomitization in the Kanye Basin can be explained only by peculiar environmental conditions at the time of deposition, or particular localized processes during the early and late diagenesis.

5.3. Chemical composition of the dolomites Sr, Na, Fe and Mn can be used to constrain dolomite evolution during diagenesis as they reflect both composition of the parent fluids and the extent of water-rock interaction during diagenesis. Under deep burial conditions Sr and Na contents tent to be very low, as these elements are strongly mobile during diagenesis, while Fe and Mn contents are generally elevated reflecting the strong reducing conditions of deep subsurface settings (Warren, 2000). On the other hand, strong depletion of Sr and Na may point toward a dolomitization process within the marine-meteoric mixing zone (Warren, 2000). Dolomite that precipitated from a mixture of meteoric water with seawater has a Sr content below 125 ppm (Warren, 2000). Ancient dolomites precipitated into hypersaline waters, on the other hand, show Sr values higher than 550 ppm (e.g., Tucker and Wright, 1990). The dolomite under investigation shows average Sr concentration below 40 ppm indicative of a parent water with anomalously low Sr contents or diagenetic loss of Sr. Ca and Sr show a positive correlation in the studied samples reflecting a likely loss of Sr during dolomitization (Veizer, 1983, Bau and Alexander, 2006; Frauenstein et al., 2009). In the studied samples Na contents are always below the XRF detection limit suggesting a low-Na precursor or diagenetic removal of Na as burial dissolution may have reduced the original contents of Na (Warren, 2000).

Fe and Mn contents tend to increase during later episodes of diagenetic recrystallization as the Fe and Mn distribution depends from the redox conditions of the mineralizing fluids (Tucker and Wright, 1990). Fe and Mn in the studied samples show average concentrations of ca. 2000 and 3500 ppm, respectively. In the Phanerozoic record the early, near-surface dolomites tend to have low Fe and Mn contents owing to oxidizing conditions, while later burial dolomites may have much higher levels due to the reducing conditions that typify deep basinal waters (Warren, 2010). For Precambrian dolomite, this trend can be opposite as the deposition environment was primarily under anoxic conditions with the result of an original enrichment of Fe and Mn at time of deposition (e.g., Veizer et al., 1992). Ferroan dolomite has been reported from partially dolomitized carbonate platforms (i.e. Latemar platform, Italy) as result of the interaction between formational water and mafic dikes (Blomme et al., 2017). The interaction between the formation water and the mafic dikes intrusion causes an overall Fe (and Mn) enrichment favoring precipitation of ferroan dolomite during the dolomitization processes. Besides a general enrichment of Mn, the studied samples show positive correlation between Fe and Mg, and XRD analysis documents the presence of ankerite, ferroan and non-ferroan dolomite (Figs. 9A, 10). The position of the main peak in the XRD spectra of the dominant carbonate phase corresponds to Fe contents ranging between ca. 800 and >3000 ppm (cf. Blomme et al., 2017). In the study of mafic dikes, the spatial distribution of ferroan dolomite (and ankerite) vs non-ferroan dolomite likely follows the general trend of the dikes, with higher Fe-content close to the intrusions (Blomme et al., 2017), although no field evidence has been identified so far. Eventually, the enrichment of Mn and Fe within non-planar dolomite with respect to planar fabrics within the Ramonnedi Formation dolomites may confirm the later stage diagenetic dolomitization related to the emplacement of the Moshaneng Dolerites. Interestingly the Mn contents are generally higher in the unaltered stromatolitic facies when compared with the facies clearly affected by metasomatism (3249 ppm and 1504 ppm, excluding sample FK_33; Table 3). This rules out the possible enrichment due to metasomatic fluids circulation (cf. Johnson et al., 2013). Enrichment of Mn in the stromatolitic carbonate can be explained by original Mn enrichment

in Archean surface water (Holland, 1984; Fischer and Knoll, 2009). Similar enrichment of Mn hosted in carbonate phases has been reported by Johnson et al. (2013) from the Transvaal Supergroup of Griqualand West Basin. In general, Archean carbonates tend to be enriched in Mn (up to 1.3%) with respect of their Phanerozoic counterparts. Johnson et al. (2013) argue that enrichment of Mn(II) in Precambrian marine sediments is a byproduct of anoxygenic photosynthesis. Manganese, in fact, has played a pivotal role in the preoxygenated Archean oceans by acting as electron donor for anoxygenic photosynthetic organisms (Johnson et al., 2013).

5.3.1 REE and trace element distributions and anomalies of ancient stromatolites The REE systematics of ancient stromatolites have previously been used to investigate the evolution of Precambrian oceans (Kamber and Webb, 2001; Van Kranendonk et al., 2003; Bolhar et al., 2015); and define the environment of formation of stromatolites (Bolhar and Van Kranendonk, 2007; Awramik and Buchheim, 2009). Microbial carbonates show a general enrichment of REE+Y with respect to inorganic marine precipitates (e.g., Webb and Kamber, 2000; Franchi et al., 2015). According to Banner et al. (1988) formation of secondary phases during dolomitisation involving non-marine fluids, followed by major textural recrystallization leaves REE patterns unaltered (see also Van Kranendonk et al., 2003). REE are considered more stable than other geochemical proxies (Van Kranendonk et al., 2003) because they replace Ca2+ in the carbonate lattice (e.g., Zhong and Mucci, 1995), and several authors have demonstrated that dolomitization does not affect the REE pattern (Bau and Alexander, 2006 and references therein). It is therefore likely that, unless total dissolution is achieved, the REE pattern of dolomite is left unaltered after major diagenetic recrystallization (Webb and Kamber, 2000; Kamber and Webb, 2001; Kamber et al., 2003). Nevertheless, other authors have demonstrated how circulation of basement fluids through carbonate sediments may alter the original Ce and Y distribution (e.g., Nothdurft et al.,

2004). It seems therefore likely that dolomite stability, with respect to the REE distribution, cannot be considered as a general rule and that dolomite recrystallization and neomorphism induced by metasomatic fluids might alter the pristine signature of marine precipitates (e.g., Land, 1992; Machel, 1997). The studied samples show that the average ∑REE content of altered facies is slightly higher than the stromatolitic dolomite. Higher ∑REE contents, often coupled with a positive Eu anomaly and lack of negative Ce anomaly, can be attributed to the mixing of hydrothermal fluids with sea water (e.g., Bau and Dulski, 1999; Kamber and Webb, 2001). The Y/Ho ratio is often used as a proxy for marine precipitates (e.g., Nozaki et al., 1997; Bau and Dulski, 1999; Franchi et al., 2015; 2017) but is also considered to be strong indicator of hydrothermal fluids circulation and REE fractionation (e.g., Klinkhammer et al., 1994; Bau and Dulski, 1999; Kamber and Webb, 2001). Previous works on Paleoproterozoic marine sediments have revealed that the input of hydrothermal fluids was characterized by a chondritic Y/Ho ratio (Bau and Dulski, 1999; Bolhar et al., 2005), whereas surface seawater and related marine precipitates were characterized by a super-chondritic Y/Ho ratio (up to 90) similar to the present-day seawater (Bau and Dulski, 1999; Franchi et al., 2017 and references therein). The Y/Ho ratio is rather uniform throughout the data set presented in this research, suggesting that metasomatic fluids did not alter original Y/Ho as both altered and unaltered facies have a near-chondritic ratio (average 32.07). The relatively uniform Y/Ho chondritic ratio may, therefore, points toward an unlikely Y depletion within the parent water or a diagenetic mobilization of Y which has been proven to be decoupled from other REE under certain conditions (e.g., Bau and Dulski, 1999; Franchi et al., 2017). Nothdurft et al. (2004) have suggested that dolomitization associated with deep fluids circulation might alter the overall Y distribution and, as a result, the Y/Ho ratio. It is therefore possible that the Y depletion, rather than being inherited from parent water, was induced by the dolomitization process itself. Kamber and Webb (2001) reported REE+Y distribution of several carbonate samples from the Campbellrand platform in South Africa. The samples were mostly

limestones and stromatolitic limestones showing a REE pattern typical of Archean oceans (La positive anomaly, Y/Ho super-chondritic ratio and no Ce anomaly; Bau and Dulski, 1996; Kamber and Webb, 2001). Interestingly, the only sample with clear evidences of partial dolomitization (sample 4 in Kamber and Webb, 2001) is the one showing a near-chondritic Y/Ho ratio and a weak La anomaly. This sample of partial dolomitized giant domal stromatolite (sample 4 in Kamber and Webb, 2001) may be analogue of the domal stromatolites of the Kanye Basin revealing a clear diagenetic effect on the distribution of La and Y related to dolomitization. Archean oceans were largely influenced by inputs of hydrothermal fluids characterized by chondritic Y/Ho ratios and a marked Eu anomaly (e.g., Michard, 1989; Klinkhammer et al., 1994; Bau and Dulski, 1999; Kamber and Webb, 2001; Franchi et al., 2017). There is consensus in the scientific community about the accuracy of the Eu anomaly as a proxy for the identification of hot fluids activity within marine precipitates as Eu is particularly enriched in high-temperature hydrothermal fluids (e.g., Klinkhammer et al., 1994; Bau and Dulski, 1999; Douville et al., 1999; Bolhar et al., 2005; Bau and Alexander, 2006; Franchi et al., 2015; 2016). Nevertheless, a strong positive Eu anomaly is lacking in most of the studied samples (Table 3). The average Eu*/Eu value in the Ramonnedi Formation dolomites is 1.10 with a maximum for the unaltered dolomite of 1.21 (FK_23). Only sample FK_33 retains a clear positive Eu anomaly ascribable to localized (i.e. within the metasomatic alteration halo due to the intrusions) precipitation of minerals from hydrothermal fluids (Table 3). This specimen was sampled less than 2 km from the asbestos mine in Moshaneng village (Fig. 2B). The presence of a strong Eu anomaly in sample FK_33 indicates circulation of hydrothermal fluids coupled with lack of a Fe-oxyhydroxide precipitation around the vents that would otherwise scavenge all the REE, obliterating all the original anomalies (Kamber and Webb, 2001). Other samples from altered facies, collected in vicinity of dolerite intrusions, do not show a consistent positive Eu anomaly (Table 3). It is therefore likely than the fractionation of Eu from hot mineralizing fluids is probably phase-related. The sub-

samples with the highest Eu anomaly, in fact, are those with evidence of dolomite replacement by calcite and Mg-silicates such as talc (Table 1). Calcite and talc are common neomorphic phases during calc-silicate metasomatism (e.g., Aleksandrov, 2011). It is likely, therefore, that the weak acid leaching has brought into solution the neomorphic calcite which is the carbonate phases bearing the extremely positive Eu anomaly (Tables 1, 3) and that, therefore, the positive Eu anomaly reflects the circulation of metasomatic fluids rather than a pristine signature of the depositional environment. Overall the samples of unaltered dolomite present a slightly negative Eu anomaly (Table 3). Precambrian sedimentary rocks often show negative Eu anomaly (e.g., McLennan and Taylor, 1991) inherited from their granitic source (Gao and Wedepohl, 1995). Similar negative Eu anomaly (in the PAAS normalized REE patterns) have been reported from the Archaean Mushandike limestone in Zimbabwe (Kamber et al., 2004). Kamber et al. (2004) consider the presence of such negative Eu anomaly as a normalization problem. Once normalized with the country rock (tonalite gneiss) the Eu anomaly disappears. This because Zimbabwean tonalite gneisses are characterized by very strong Eu depletion. Similar consideration can be done in the study area where the country rocks were mostly Archean granites (i.e. Gaborone Granites). There are no available REE data on the geochemical composition of the Gaborone Granites in Botswana, nevertheless it is reasonable to assume they were characterized by negative Eu anomaly, a distinctive feature of most Archean granites (e.g., Gao and Wedepohl, 1995). Moreover, it is likely to assume that during the Archean along the Ramonnedi Formation platform, in a closed epicontinental sea, not influenced by input of Eu-rich hydrothermal water, the seawater was characterized by an overall Eu depletion (e.g., Kamber et al., 2004). This would explain the consistent negative Eu anomaly in the PAAS-normalized REE patterns of unaltered dolomite from the Taupone Formation. The lack of a Ce anomaly can be considered to be evidence for precipitation under reducing conditions (Kamber and Webb, 2001). In oxygenated environments Ce(III) is preferentially

oxidized to its more stable form Ce(IV) and easily scavenged by Fe-oxyhydroxides resulting in an overall negative Ce anomaly in seawater and seawater precipitates (e.g., Hu et al., 2014). Carbonates precipitated from anoxic water should therefore record a lack of strongly negative Ce anomaly (e.g., Van Kranendonk et al., 2003). The studied samples reveal a general lack of marked positive or negative Ce anomaly with average values of ca. 0.91 (Table 3). The overall lack of a negative Ce anomaly from the samples suggests that during deposition of the Ramonnedi Formation dolomite the pO2 of the surface seawater was too low to allow the oxidation of Ce(III) (e.g., Bau and Alexander, 2006). The lack of a Ce anomaly and consequent reducing conditions of the surface water at the time of the Ramonnedi Formation dolomite formation is not surprising considering that this unit was deposited before the BIF of the Masoke Formation (Fig. 3) and reflects depositional conditions characterized by anoxic seawater and a reducing Archean atmosphere (e.g., Kato et al., 1998; Van Kranendonk et al., 2003). Nevertheless, in the Ce/Ce* vs Pr/Pr* plot (Fig. 9B-C) the samples cluster in three distinct groups: i) sample FK_33 shows a slightly negative Pr anomaly and a Ce anomaly close to unity; ii) other altered samples (i.e. FK_28 and FK_29) show a slightly positive Pr anomaly and a slightly negative Ce anomaly; iii) samples of unaltered dolomite show a decrease of the Ce/Ce* values where the Pr/Pr* values increase. The trend in figure 9C confirms that the negative Ce anomaly recorded by 2 dolomite samples (FK_8B and FK_32A) is real and might reflect oxygenated seawater while most of the samples of stromatolitic dolomite record either a positive La anomaly and lack of a real negative Ce anomaly. The fluctuation of Ce anomaly among stromatolitic samples (and often within the same sample; Table 3) reflects changes in oxygen content in the water directly in contact with the stromatolites. Assuming that the dolomitization processes did not affect overall Ce distribution, the negative Ce anomaly recorded in specific lithofacies (i.e. domal and micro-columnar stromatolites) may reflect the presence of oxygen hot spots or ‘oxygen oases’ (Kasting et al., 1992; Bau and Alexander, 2006; Riding et al., 2014) probably related to the metabolic activity of phototrophs. Waldbauer et al. (2009) have demonstrated that photosynthesis and other metabolic pathways have arisen early in the Archean (see

also Des Marais, 2000). In an overall anoxic Archean ocean the presence of thriving communities of stromatolites, allegedly produced by colonies of phototrophs, might create hot spots for the diffusion of O2 as a byproduct of their metabolism (oxygen oases). The presence of such hot spots of oxygenated water may represent the first stage of the gradual increase of oxygen content within otherwise anoxic oceans that culminated during the GOE.

6. CONCLUSIONS The stratigraphic, petrographic and geochemical characterization of dolostone from the Taupone Dolomite Group led to the following conclusions: 1. The sedimentary sequence of the Lower Transvaal Supergroup preserved in the Kanye Basin records a deepening upward trend. The Taupone Dolomite Group sequence begins with a tidal channel fill sequence, grades into a well-developed carbonate platform and is capped by oolitic units and by banded chert probably deposited under subtidal conditions. The bottom of the Ramonnedi Formation represents the first stage of transgression with transition from a shallow marine environment (between fair wave base and high tide level) to a low lying tidal mud flat and to a protected lagoon characterized by evaporative processes, while the chert (and subsequent BIF of the Masoke Formation) marks the drowning of the carbonate platform. 2. The Ramonnedi Formation dolomite clearly represents a diagenetic product as shown by the crystal fabrics, dominated by non-planar xenotopic and hypidiotopic dolomite. The paucity of pristine microcrystalline dolomite, together with the dominance of coarse grained dolomite, is indicative of recrystallization (dolomitization) of the original platform carbonates. The presence of both idiotopic planar and xenotopic non-planar dolomite within most of the studied dolostone samples suggests multiple-stage diagenesis.

The petrography of the Ramonnedi Formation dolomite reflects an early stage dolomitization during which dolomite precipitated within the marine/meteoric mixing zone under shallow burial conditions characterized by an overall Mn and Fe enrichment typical of Archean carbonates. A late stage of dolomitization has probably occurred at deepest burial condition inducing recrystallization of microcrystalline and idiotopic dolomite to xenotopic dolomite. This change occurred under reducing condition and left unaltered the original Fe and Mn enrichment. 3. The circulation of metasomatic fluids through the Ramonnedi Formation sediments have induced the replacement of dolomite by talc, chrysotile and other alteration minerals localized close to the intrusion in the Moshaneng village. The imprint by metasomatism on the rest of the Ramonnedi Formation dolomites is not clear and its spatial distribution is not known. A better understanding of metasomatic processes in the area will increase our knowledge of the distribution of metasomatic processes and mineralization. 4. Sr and Na have been diagenetically removed from the carbonates, probably during the early stages of dolomitization. On the other hand, the high Fe and Mn contents reflect an original enrichment of these elements in the anoxic Archean marine water. Archean carbonates tend to be enriched in Mn with respect of their Phanerozoic counterparts. This Mn-enrichment in the stromatolitic carbonates of the Ramonnedi Formation is probably byproduct of anoxygenic photosynthesis. Manganese, in fact, has played a pivotal role in the pre-oxygenated Archean oceans by acting as electron donor for anoxygenic photosynthetic organisms. 5. Although REE are normally considered robust geochemical proxies for the investigation of ancient carbonates, the effects of dolomitization and diagenesis must be taken into account. The REE+Y distribution of the Ramonnedi Formation dolomites have been clearly affected by the dolomitization processes. Particularly the process of dolomitization of the carbonate platform has led to a chondritic Y/Ho ratio and to a weakening of the La positive anomaly

that normally characterize Archean marine precipitates (Kamber and Webb, 2001). It is therefore possible that Y, with respect to other REE, is more prone to be mobilized during deep burial dolomitization leading to a shift of Y/Ho ratio toward near chondritic values. 6. The overall lack or the presence of a negative Eu anomaly suggest that the basin where these sediments precipitated was completely restricted from open seawater. A general lack of a Ce anomaly points toward a dominant reducing conditions of the surface water at the time of the Ramonnedi Formation dolomite formation. This is not surprising considering that this unit was deposed in anoxic seawater and reducing Archean atmosphere. Nevertheless, the extent to which Ce distribution has been affected by the dolomitization is not known. It is worth to note, anyway, that the data set shows a broad range of the Ce anomaly, within the 0.27-1.08 range, among stromatolitic samples probably reflecting the original oxygen content variability. 8. Besides the complications induced by multiple stage of diagenetic processes, it appears reasonable that the stromatolites thriving within a closed, epeiric sea became the hot spots for the production and diffusion of oxygen within the anoxic Neoarchean/Paleoproterozoic seawater ca. 0.3-02 Ga before the onset of the Great Oxidation Event (GOE). The presence of such hot spots of oxygen production within the stromatolitic dolostone from the Lower Transvaal Supergroup may represent the starting point of the gradual increase of oxygen content within otherwise anoxic oceans. After adequate consideration of post-depositional processes, the Ramonnedi Formation stromatolites are a valuable proxy for the investigation of the transition from an Archean inhospitable anoxic ocean to more oxygenated waters that culminated during the GOE. 9. Although the common beliefs consider REE as stable during multiple stage diagenesis/dolomitization, the study of Neoarchean dolomites from the Ramonnedi Formation has demonstrated that overall REE+Y pattern of ancient carbonates can indeed be altered by diagenesis, particularly by late dolomitization under burial conditions.

ACKNOWLEDGEMENTS This research was funded by BIUST Initiation Grant 2016 to FF. Thanks are due to R. Tisane for her support during the preparation of the samples, to all the BIUST 3RD year Geology students 2015/2016 for the exciting days in the field, and to G. Angus for the thorough proof reading of the manuscript. The author thanks R. Bolhard and R.B.M. Mapeo for the helpful comments that have contributed to greatly improve this manuscript.

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Sumner, D.Y., 1997. Carbonate precipitation and oxygen stratification in late Archean seawater as deduced from facies and stratigraphy of the Gamohaan and Frisco formations, Transvaal Supergroup, South Africa. Amer. J. Sci. 297, 455-487. Sumner, D.Y., Grotzinger, J.P., 2004. Implications for Neoarchaean ocean chemistry from primary carbonate mineralogy of the Campbellrand-Malmani Platform, South Africa. Sedimentology 51, 1-27. Sumner, D.Y., Beukes, N.J., 2006. Sequence Stratigraphic Development of the Neoarchean Transvaal carbonate platform, Kaapvaal Craton, South Africa. S. Afr. J. Geol. 109, 11-22. Tang, L., Santosh, M., Tsunogae, T., Maruoka, T.,2016. Paleoproterozoic meta-carbonates from the central segment of the Trans-North China Orogen: Zircon U–Pb geochronology, geochemistry, and carbon and oxygen isotopes. Precambrian Res 284, 14-29. Taylor, S.R., McLennan, S.M., 1985. The Continental Crust: Its Composition and Evolution. Blackwell Scientific Pub, Palo Alto (CA) 328 pp. Tucker, M., Wright, V.P., 1990. Carbonate Sedimentology. Blackwell Scientific Publications, Oxford, 482 pp. Van Kranendonk, M.J., Webb, G.E., Kamber, B.S., 2003. New geological and trace element evidence from 3.45 Ga stromatolitic carbonates in the Pilbara Craton: support of a marine, biogenic origin and for a reducing Archaean ocean. Geobiology 1, 91-108. Van Tuyl, F.M., 1918. Depth of dolomitization. Science 48, 350-352. Veizer, J., 1983. Trace elements and isotopes in sedimentary carbonates. In: E., Roedder (Ed.) Carbonates: Mineralogy and Chemistry. Rev. Mineral. 11, 265-300. Veizer, J., Clayton, R.N., Hinton, R.W., 1992. Geochemistry of Precambrian carbonates: IV. Early Paleoproterozoic (2.25 ± 0.25 Ga) seawater. Geochim. Cosmochim. Acta 56, 875-885.

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TABLE CAPTIONS Table 1. Sample location, lithofacies and mineralogical composition (from XRD). Table 2. Major and trace elements distribution from the studied samples. Table 3. Ca/Mg ratio, ∑REE and REE anomalies from studied samples.

FIGURE CAPTIONS Figure 1. Schematic map of southern Africa showing the distribution of the Transvaal Supergroup rocks, grey broken line indicates the extension of the Kaapvaal craton (modified from Eriksson et al., 2002). Kanye is indicated by a black star.

Figure 2. Geological map of the great Kanye area (A). B) Schematic geological map of the Moshaneng village area (north west from Kanye village). See box in A. C) Schematic geological map of the Diphawana hills area (south east from Kanye village). See box in A. Modified from the 1998 digital edition of the Geological map of Botswana (Key and Ayres, 2000).

Figure 3. Lithostratigraphic chart of the Lobatse Volcanic Group and Transvaal Supergroup in Botswana (modified from Aldiss et al., 1989; Catuneanu and Eriksson, 1999; Mapeo et al., 2006).

Figure 4. Lithostratigraphic log of the Lower Transvaal Supergroup in the Diphawana hills area (Fig. 2B).

Figure 5. Outcrop views of the Ramonnedi Formation. A) massive dolostone (b; FK_5B) interbedded with murdock (a, d) and siltstone with mud-cracks (c); B) domal stromatolites (FK_8A); C) stratiform stromatolite (FK_8B); D) stratiform stromatolite covered by microcolumnar stromatolite (FK_8D); E) large scale domal stromatolites (FK_8E); F) columnar stromatolite (FK_8F); G) Upper unit, dolomite and chert interbedding with oolitic chert (FK_27) levels below the knife (H); H) Oolitic chert (knife is ca. 8 cm); I) Cherty dolostone with “ribbon” fabrics (FK_28).

Figure 6. Photomicrographs of the dolomite microfacies. A) Xenotopic mosaic of coarse non-planar dolomite (FK_32); B) Hypidiotopic mosaic of planar subehedral dolomite and microcrystalline dolomite (FK_30); C) Idiotopic mosaic of planar dolomite with evidences of dedolomitization in the center of dolomite rhombohedra (FK_8A); D) Rhombohedra of dolomite within microcrystalline dolomite matrix (mdo) and microcrystalline replacive quartz (mqz) (FK_30); E) Laminae of microcrystalline quartz cutting across hypidiotopic mosaic of subhedral dolomite (FK_17); F) Microcrystalline quartz and replacive syntaxial cement (FK_8E); G-H) Oolitic chert (FK_18). I) FE-SEM photomicrographs of sample FK_8E showing spherules engulfed within dolomite crystals (field of view 70 µm; 15 kV, 10 mm working distance). A-C, G: planar polarized light. D-F, H: cross polarized light.

Figure 7. PAAS-normalized REE patterns of the studied samples: A) Massive dolostone. B) Domal and columnar stromatolites showing homogeneous flat REE patterns. C) Stratiform stromatolites. D) Micro-columnar stromatolites. E) Altered facies (dotted line refers to sample FK_32B); FK_33 sub-samples show partial REE pattern. F) Oolitic chert.

Figure 8. Schematic lithostratigraphic column of Ramonnedi Fm. in the two study localities, Diphawana Hills and Moshaneng Village (Fig. 2), showing the variations of REE patterns with lithofacies: massive dolomite and mudrock (A), domal stromatolite(B), stratiform and micro-columnar stromatolite (C), large scale domal and columnar stromatolite (D), oolitic chert and dolostone (E), ‘ribbon dolomite’ (F). Refer to figure 4 for keys.

Figure 9. Cross plots of selected trace elements and anomalies: A) Mg vs Fe plot showing a good correlation excluding outliers from mudrocks (FK_03 and FK_5B_A). B) Cross plot of Ce and Pr anomalies from altered facies. C) Cross plot of Ce and Pr anomalies from stromatolitic dolomites (fields were modified from Bau and Dulski, 1996).

Figure 10. Cross plot of the position of main XRD peaks (d104) of dolomite/ankerite (see Table 1) and Fe content from ICP-MS analysis.

Supplementary Material S1. Hand specimens of dolostone and stromatolites. Solid diamonds show sites of subsampling for XRF and ICP-MS analyses (withe diamonds) and for XRD analyses (red diamonds). Where indications of subsampling are lacking the data refer to the whole rock analysis.

Supplementary Material S2. Hand specimens of altered dolostone. Solid diamonds show sites of subsampling for XRF and ICP-MS analyses (withe diamonds) and for XRD analyses (red diamonds). Where indications of subsampling are lacking the data refer to the whole rock analysis.

Location Lat. Long.

Lithofacies

Mineralogica

Massive dolostone

Dolomite

Massive dolostone

Dolomite, Quartz

Mudrock

Dolomite, quartz, muscovite, m

E25.33005°

Domal stromatolite

Dolomite, Quartz

S25.09954°

E25.33005°

Domal stromatolite

Dolomite, Quartz

FK_08C

S25.09954°

E25.33005°

Dolomite, Quartz

FK_08D

S25.09954°

E25.33005°

Large domal stromatolite Stratiform to micro-columnar strom. Large domal stromatolite

Dolomite, Quartz

Large domal stromatolite

Dolomite, Quartz

FK_08E C

Large domal stromatolite

Dolomite, Quartz

FK_08F A

Domal stromatolite

Dolomite

FK_08F B

Domal stromatolite

Dolomite

FK_08F C

Domal stromatolite

Dolomite

FK_08F D

Domal stromatolite

Ankerite/Dolomite

Domal stromatolite

Dolomite

FK_08F F

Domal stromatolite

Dolomite

FK_08F G

Domal stromatolite

Dolomite

FK_08F H

Domal stromatolite

Dolomite

FK_08F I

Domal stromatolite

Ankerite/Dolomite, Quartz

FK_17 A

Stratiform stromatolite

Dolomite, Quartz

Stratiform stromatolite

Dolomite, Quartz

Stratiform stromatolite

Dolomite, Quartz

FK_03

S25.09568°

E25.32494°

FK_05B A

S25.09597°

E25.32555°

FK_08A

S25.09954°

FK_08B

FK_05B B

FK_08E A FK_08E B

FK_08F E

FK_17 B

S25.09995°

S25.09995°

S25.09932°

E25.33010°

E25.33010°

E25.32988°

FK_17 C

Dolomite, Quartz

FK_18

S25.10201°

E25.32342°

Oolitic dolostone

Dolomite, Quartz

FK_23

S25.07360°

E25.28564°

Massive dolostone

Dolomite, Quartz

Oolitic to layered dolostone

Dolomite

Oolitic to layered dolostone

Dolomite, Quartz

FK_27 C

Oolitic to layered dolostone

Dolomite

FK_27 D

Oolitic to layered dolostone

Dolomite, Quartz

Layered metasomatic dolostone*

Calcite, Chrysotile, Ankerite/Do

Layered metasomatic dolostone*

Calcite, Chrysotile, Forsterite, C

Layered metasomatic dolostone*

Calcite, Chrysotile, Dolomite

Layered metasomatic dolostone*

Calcite, Chrysotile, Dolomite

FK_29 B

Layered metasomatic dolostone*

Tremolite, Chrysotile, Dolomite

FK_30 A

Micro-columnar stromatolite

Dolomite, Quartz

Micro-columnar stromatolite

Dolomite, Quartz

Micro-columnar stromatolite

Dolomite, Quartz

FK_27 A FK_27 B

S25.08876°

E25.30728°

FK_28 A FK_28 B

S25.081179°

E25.283460°

FK_28 C FK_29 A

FK_30 B FK_30 C

S24.90873°

S25.10614°

E25.23376°

E25.32395°

FK_30 D

Micro-columnar stromatolite

Dolomite, Quartz

FK_30 E

Micro-columnar stromatolite

Dolomite, Quartz

FK_30 F

Micro-columnar stromatolite

Dolomite, Quartz

FK_31 A

Layered metasomatic dolostone*

Calcite, Chrysotile, Dolomite

Layered metasomatic dolostone*

Calcite, Chrysotile, Dolomite, F

FK_31 D

Layered metasomatic dolostone*

Calcite, Chrysotile, Dolomite/A

FK_32 A

Massive dolostone

Dolomite, Quartz

Vein

Dolomite, Quartz

FK_32 C

Massive dolostone

Dolomite

FK_32 D

Massive dolostone

Dolomite

FK_33 A

Layered metasomatic dolostone*

Talc, Diopside, Tremolite (?)

Layered metasomatic dolostone*

Calcite, Talc, Diopside, Tremol

FK_33 C

Layered metasomatic dolostone*

Calcite, Talc, Diopside, Tremol

FK_33 D

Layered metasomatic dolostone*

Talc, Diopside, Tremolite (?)

FK_31 C

FK_32 B

FK_33 B

S24.85019°

S24.92481°

S24.90881°

E25.25147°

E25.25705°

E25.22580°

* Former ‘ribbon dolomite’ (Aldiss et al., 1989)

Table 1. Sample location, lithofacies and mineralogical composition (from XRD).

FK_ 03

FK_0 5B A

FK_0 5B B

FK_0 8A

FK_0 8B

FK_0 8C

FK_0 8D

FK_0 8E A

FK_0 8E C

FK_0 8E B

Mg O K2 O Ca O TiO 2

Mn O Fe2 O3 Al2 O3 SiO 2

P2O 5

LOI

FK_0 8F A 23.6 6 0.19 31.8 9

FK_0 8F B 24.0 7 0.16 32.0 8

FK_0 8F C 22.8 7 0.29 30.4 4

0.02

0.02

0.02

0.88

0.86

0.82

0.60

0.60

0.57

0.50

0.51

0.60

2.02

1.80

4.23

0.01 44.7 9

0.01 44.1 6

0.01 44.2 6

FK_0 8F D

FK_0 8F E

FK_0 8F F

Mg

0.0 02 0.4 40 621 00

0.00 3 1.17 0 6400 0

0.00 4 1.53 0 1130 0

0.00 2 0.13 0 1020 00

0.00 1 0.18 0 573 0

0.00 1 0.07 0 8080 0

0.00 1 0.14 0 6160 0

0.00 1 0.09 0 6710 0

0.00 1 0.09 0 7710 0

0.00 1 0.07 0 7860 0

0.00 2 0.12 0 9830 0

0.00 2 0.13 0 1000 00

0.00 2 0.31 0 1260 00

0.00 3 0.22 0 1550 00

0.00 1 0.12 0 1010 00

0.00 2 0.15 0 8670 0

Al

169

472

4650

229

26

270

325

88

116

114

247

230

99

318

91

227

K

200 ### ## 0.2 00 2.3 00 0.9 70 1.4 00 373 0 700 0 0.2 58 ### ## 0.4 00 9.0 00 0.9 00 0.1 07 0.2 80 0.7 00 5.0 00 0.3

860 1290 00 2.70 0 2.70 0 0.72 0 3.90 0

5950 2460 0 1.70 0 1.10 0 5.90 0 17.3 00

300 1690 00 0.80 0 3.90 0 2.36 0 1.90 0

30 910 0 3.30 0

410 1040 00 0.60 0 2.70 0 1.99 0 2.00 0

130 1090 00

160 1270 00 2.40 0 0.90 0 1.30 0

220 1630 00 0.40 0 4.60 0 1.68 0 1.50 0

190 1650 00 0.30 0 3.00 0 1.68 0 1.20 0

1.50 0 1.64 0 0.90 0

340 2700 00 0.50 0 5.90 0 2.63 0 2.30 0

40 1800 00

2.70 0 0.80 0 1.20 0

150 1300 00 0.20 0 2.70 0 0.99 0 1.60 0

2120 00

0.13 0 0.30 0

320 1380 00 0.70 0 3.00 0 2.85 0 1.70 0

2.70 0 1.05 0 0.50 0

190 1520 00 0.40 0 3.20 0 1.65 0 1.50 0

4370

535

5180

313

4210

2370

2030

2410

2370

5250

5600

6700

8020

5360

4830

8260 1.01 0 179. 000 3.30 0

2580 2.24 0 333. 000 7.40 0 6.00 0 1.31 0 0.15 4 0.31 0 0.90 0 4.50 0 8.82

2830 0.37 4 6.23 0 1.40 0 7.00 0 0.77 0 0.07 9 0.33 0

159 0.04 8 3.73 0 3.80 0 12.0 00 0.11 0 0.01 9 1.83 0 0.70 0 1.40 0 0.04

2470 0.89 9 59.9 00 2.10 0 3.00 0 0.88 0 0.08 3 0.33 0

2460 0.62 7 37.1 00 0.60 0 5.00 0 1.69 0 0.15 5 0.49 0

2010 0.56 9 580. 000 0.90 0 4.00 0 0.51 0 0.04 6 0.15 0

2280 0.47 0 211. 000 0.50 0 3.00 0 0.55 0 0.05 0 0.19 0

2140 0.31 5 23.9 00 0.10 0

3270 0.46 6 33.7 00 0.30 0 4.00 0 0.82 0 0.07 9 0.25 0

3630 0.26 9 26.2 00

0.51 0 0.05 0 0.13 0

3040 0.40 6 28.8 00 0.30 0 5.00 0 0.76 0 0.07 0 0.21 0

2850 0.30 1 31.2 00 1.50 0 2.00 0 0.72 0 0.05 3 0.22 0

2580 0.35 7 29.0 00 0.20 0 3.00 0 0.70 0 0.06 6 0.24 0

2.20 0 0.30

4.20 0 0.45

2.20 0 0.13

2.70 0 0.13

2.70 0 0.13

4.80 0 0.22

5.90 0 0.20

19.2 00 -

4490 0.88 9 324. 000 0.80 0 8.00 0 0.97 0 0.08 6 0.29 0 0.70 0 8.40 0 0.24

3.60 0 0.04

3.80 0 0.16

Li Be

Ca Sc Ti V Cr Mn Fe Co Ni Cu Zn Ga Ge As Se Br Rb

1.83 0 0.25 0 0.63 0 2.50 0 6.30 0 0.99

2.70 0 0.25

14.0 00 0.92 0 0.07 9 0.28 0

50 ### ## 3.1 90 0.1 70

0 14.4 00 8.72 0 0.18 0

0 6.30 0 3.67 0 0.14 0

0 13.2 00 1.21 0 0.16 0

0 1.00 0 0.10 1 0.10 0

0 9.40 0 1.40 0 0.09 0

0 10.1 00 2.77 0

0 7.80 0 0.72 3

0 8.40 0 0.81 3

0 9.10 0 0.73 4

0 12.3 00 1.04 0

-

-

-

-

Nb

-

-

-

-

-

-

-

-

-

0.01 6

0.01 0 0.00 3

-

0.0 11 0.0 30 0.0 15

0.04 0 0.01 1 0.03 0 0.04 0

50.4 00 0.03 0 0.02 9 0.07 0 0.07 9

-

Mo

11.6 00 0.01 0 0.00 8

0.00 5

Sr Y Zr

Cd In

Hf

1.0 00 1.2 30 2.7 90 0.3 59 1.5 60 0.3 90 0.0 64 0.4 43 0.0 80 0.4 37 0.0 86 0.2 23 0.0 28 0.1 89 0.0 30 0.0 04

0.00 8 0.20 0 0.00 2 8.00 0 2.10 0 5.25 0 0.77 9 3.52 0 1.11 0 0.19 4 1.20 0 0.23 5 1.41 0 0.28 5 0.85 3 0.12 0 0.79 1 0.11 6 0.00 9

Ta

-

-

W

-

-

Re

-

Pb

### ##

0.00 1 10.7 00

Bi

-

-

Sn Sb I Cs Ba La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu

Tl

-

0.02 3 14.0 00 2.01 0 2.75 0 0.46 3 1.89 0 0.48 3 0.08 6 0.52 7 0.09 5 0.54 4 0.11 3 0.33 1 0.04 5 0.26 2 0.04 3 0.00 5 0.00 2 0.00 3 0.02 5 7.01 0 0.01 0

0.01 7 2.00 0 1.30 0 2.50 0 0.29 2 0.99 2 0.20 3 0.04 1 0.21 3 0.03 6 0.18 6 0.03 8 0.11 4 0.02 4 0.09 9 0.02 2 0.02 0 3.19 0 0.02 0

0.00 4 0.31 1 0.30 3 0.05 7 0.10 0 0.02 4 0.01 1 0.24 0 0.02 5 0.02 2 0.01 0 0.01 6 0.04 2 0.02 5 0.03 5 0.00 3 0.00 2 0.11 0 0.00 4 0.02 4 16.1 00 0.17 0

12.3 00 1.12 0 0.30 0

0 17.2 00 1.56 0 0.10 0

0 9.90 0 0.94 1

0 9.40 0 0.94 4

-

0 12.9 00 1.18 0 0.10 0

-

-

-

-

-

-

-

-

-

-

-

-

-

0.09 0

1.14 0

0.00 2

0.00 2

0.00 2

0.00 4

0.00 7

-

0.00 3

0.00 3

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

0.01 4 0.30 0

-

-

0.00 5

-

-

0.01 1

0.05 0 0.02 0

0.00 8

-

-

2.00 0 1.34 0 2.79 0 0.33 7 1.23 0 0.25 0 0.04 8 0.21 7 0.04 1 0.22 4 0.04 5 0.13 4 0.01 9 0.11 9 0.01 9

1.00 0 0.82 4 1.59 0 0.18 3 0.67 0 0.12 1 0.02 5 0.11 1 0.01 8 0.10 8 0.02 3 0.06 5 0.00 9 0.05 1 0.00 9

2.00 0 0.93 8 1.77 0 0.20 3 0.72 3 0.13 9 0.02 7 0.12 6 0.02 1 0.11 8 0.02 3 0.07 2 0.00 9 0.05 8 0.00 9

-

1.61 0 2.84 0 0.28 4 1.06 0 0.21 3 0.04 6 0.19 8 0.03 3 0.18 7 0.03 5 0.11 0 0.01 7 0.10 6 0.01 8

2.00 0 1.97 0 3.68 0 0.40 1 1.42 0 0.29 1 0.06 1 0.27 4 0.04 2 0.25 2 0.05 1 0.15 0 0.02 1 0.13 2 0.02 0

2.00 0 1.16 0 2.29 0 0.25 4 0.88 7 0.17 0 0.03 9 0.15 2 0.02 7 0.14 6 0.03 1 0.08 9 0.01 3 0.07 7 0.01 2

1.00 0 1.22 0 2.29 0 0.25 2 0.93 3 0.17 4 0.03 7 0.16 6 0.02 8 0.14 7 0.03 1 0.08 4 0.01 3 0.08 0 0.01 2

-

-

-

-

2.00 0 1.39 0 2.64 0 0.30 4 1.11 0 0.22 7 0.05 4 0.18 8 0.03 5 0.19 4 0.03 9 0.10 7 0.01 7 0.09 7 0.01 5 0.00 3

-

0.99 3 1.77 0 0.19 9 0.72 8 0.12 8 0.02 7 0.11 6 0.02 1 0.11 4 0.02 3 0.06 8 0.00 9 0.05 8 0.00 9

3.00 0 1.20 0 2.40 0 0.28 1 1.06 0 0.19 5 0.04 8 0.17 8 0.03 1 0.17 1 0.03 5 0.10 6 0.01 3 0.08 8 0.01 4

-

2.00 0 2.63 0 5.53 0 0.64 6 2.44 0 0.50 0 0.08 6 0.45 8 0.08 0 0.44 9 0.09 4 0.28 0 0.03 9 0.25 4 0.03 8 0.00 4

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

-

3.87 0

1.94 0

1.42 0

1.36 0

1.50 0

3.70 0

3.58 0

0.00 6 6.59 0

0.00 3 3.41 0

0.00 1 3.18 0

0.00 1 2.47 0

-

-

-

-

-

-

-

-

-

-

-

-

Th U

0.0 64 0.0 31

0.31 0 0.12 5

0.61 9 0.09 1

0.20 6 0.04 2

0.03 0 0.08 4

0.16 9 0.04 0

0.26 9 0.12 3

0.05 2 0.01 8

0.05 6 0.01 7

0.06 9 0.02 2

0.15 3 0.02 9

0.18 4 0.02 5

Table 2. Major and trace elements distribution from the studied samples.

0.09 9 0.02 7

0.18 3 0.03 7

0.04 3 0.01 7

0.13 5 0.02 4

FK_03 FK_05B A FK_05B B FK_08A FK_08B FK_08C FK_08D FK_08E A FK_08E C FK_08E B FK_08F A FK_08F B FK_08F C FK_08F D FK_08F E FK_08F F FK_08F G FK_08F H FK_08F I FK_17 A FK_17 B FK_17 C FK_18 FK_23 FK_27 A FK_27 B FK_27 C FK_27 D FK_28 A FK_28 B FK_28 C FK_29 A FK_29 B FK_30 A FK_30 B FK_30 C FK_30 D FK_30 E FK_30 F FK_31 A FK_31 C FK_31 D FK_32 A FK_32 B FK_32 C FK_32 D FK_33 A FK_33 B FK_33 C FK_33 D AVERAGE Av. Dolomite Av. Altered

Ca/Mg 1.67 2.02 2.18 1.66 1.59 1.71 1.69 1.62 1.65 1.65 1.66 1.65 1.68 1.74 1.78 1.75 1.76 1.83 1.79 1.81 1.80 1.80 1.79 1.87 1.76 1.73 1.71 1.80 2.84 4.37 11.35 10.73 9.32 1.76 1.77 1.82 1.80 1.69 1.80 2.21 8.32 3.93 7.88 4.56 2.51 1.93 44.82 8.02 2.38 3.91

La*/La 1.08 0.96 1.15 0.84 0.61 0.84 0.91 0.95 0.93 1.05 0.95 0.96 1.24 0.98 0.89 1.04 0.93 1.01 0.93 0.90 0.83 0.64 0.94 1.25 0.89 0.79 0.92 1.11 0.97 0.94 1.00 1.28 1.09 0.73 0.85 0.79 0.69 0.83 0.79 0.82 1.01 0.91 0.99 0.99 0.93 1.11 1.77 1.66 1.57 1.75

Ce*/Ce 0.98 0.88 0.70 0.84 0.27 0.87 0.93 0.92 0.90 0.94 0.93 0.92 1.08 0.94 0.91 0.97 0.95 0.95 0.91 0.92 0.87 0.83 0.90 0.87 0.91 0.91 0.94 0.97 0.95 0.92 0.96 0.95 0.94 0.86 0.94 0.95 0.87 1.07 0.94 0.94 0.95 0.97 0.41 0.95 0.90 0.98 1.01 0.96 1.06 1.00

Eu*/Eu 0.74 0.78 0.81 0.96 1.20 0.93 0.84 1.04 0.97 1.03 1.21 1.18 1.07 1.06 1.13 1.04 0.97 0.97 0.89 0.91 0.83 0.86 0.92 1.21 0.91 0.81 0.86 1.04 0.99 0.67 0.67 1.09 1.28 0.73 0.71 0.71 0.73 0.75 0.79 0.98 0.62 0.71 0.76 1.10 0.56 0.52 5.66* 3.97 1.72 1.97

Gd*/Gd 1.04 0.97 1.03 1.06 3.11 0.92 0.99 1.04 1.02 0.97 0.99 0.92 1.03 1.09 0.97 1.03 1.00 0.99 1.03 0.96 0.96 0.94 0.93 1.23 1.14 1.01 1.00 1.07 1.03 1.14 1.00 0.97 1.00 0.93 0.98 0.96 0.92 0.97 0.96 1.03 1.00 0.99 1.05 0.98 1.05 1.07 -

Pr/Yb 0.61 0.31 0.56 0.94 0.72 0.90 0.81 1.15 1.12 1.10 1.02 1.00 0.86 0.97 1.05 1.01 1.04 1.09 1.10 0.88 0.95 0.85 0.97 0.49 1.18 1.20 1.27 1.04 2.07 1.13 1.01 0.70 0.97 0.66 0.83 0.74 0.80 1.00 0.82 1.22 1.45 1.13 0.83 1.48 0.61 0.64 32.69 9.14 5.74

Gd/Yb 1.42 0.92 1.22 1.30 5.81 1.10 1.09 1.32 1.31 1.21 1.22 1.17 1.13 1.26 1.19 1.26 1.27 1.24 1.38 1.08 1.30 1.19 1.21 0.70 1.35 1.41 1.43 1.32 1.77 1.20 1.01 0.90 1.25 0.90 1.04 0.96 1.04 1.20 1.06 1.41 1.20 1.19 1.45 1.35 1.61 1.53 5.96 2.59 2.22

Y/Ho 37.18 30.60 32.48 32.27 10.00 31.18 29.50 32.13 34.74 31.64 30.06 30.41 31.91 30.59 30.45 30.95 30.06 30.19 31.04 30.16 29.32 27.99 32.88 76.32 35.43 32.66 31.12 31.28 30.60 27.94 28.78 27.99 29.09 25.69 28.29 24.90 27.17 24.94 27.48 31.03 29.12 30.09 37.82 31.30 38.23 36.00 41.78 48.75 40.00

TOT REE 7.91 17.96 9.64 6.06 1.22 6.81 13.52 3.81 4.24 4.26 5.82 6.42 6.76 8.76 5.35 5.47 6.03 6.16 6.41 11.18 8.73 8.31 1.03 1.07 3.25 2.28 4.60 0.84 17.00 15.43 37.75 5.38 5.96 13.76 16.46 7.19 6.07 5.50 10.68 12.41 12.26 17.54 18.76 10.24 14.42 11.78 99.56 14.88 0.95 1.48

3.78

1.00

0.91

1.00

1.05

1.87

1.46

32.07

10.59

2.02 9.35

0.93 1.23

0.89 0.97

0.91 1.33

1.06 1.02

0.91 5.20

1.34 1.88

31.75 33.20

7.60 20.05

* Eu/Eu* = EuSN/ (SmSN)*(GdSN)

1/2

from Taylor and McLennan (1985).

Table 3. Ca/Mg ratio, ∑REE and REE anomalies from studied samples.

Petrographic and geochemical characterization of the Lower Transvaal Supergroup stromatolitic dolostones (Kanye Basin, Botswana)

HIGHLIGHTS      

Neoarchean/Paleoproterozoic stromatolitic dolostones of the Ramonnedi Fm have been studied The sequence reveals a deepening upward trend Petrographic characterization reveals multiple stages of dolomitization Y/Ho ratios have been altered by the deep burial dolomitization Geochemical composition of the stromatolites reveals deposition in closed basin No evidence of hydrothermal input at time of deposition has been detected