Gondwana Research 17 (2010) 653–661

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Gondwana Research j o u r n a l h o m e p a g e : w w w. e l s e v i e r. c o m / l o c a t e / g r

Veracity of Neoproterozoic negative C-isotope values: The termination of the Shuram negative excursion Erwan Le Guerroué a,⁎, Andrea Cozzi b a b

Geosciences-Rennes, 263 Avenue du General Leclerc. CS 74205-35042 Rennes Cedex, France Eni Angola, Rua Nicolau Gomes Spencer, 140 P.O. Box 1289, Luanda, Angola

a r t i c l e

i n f o

Article history: Received 10 June 2009 Received in revised form 26 October 2009 Accepted 9 November 2009 Available online 17 November 2009 Keywords: Neoproterozoic Ediacaran Shuram excursion Negative carbon isotope excursion Carbon cycle

a b s t r a c t There is widespread interest in the Neoproterozoic period of the Earth's history (1000 to 542 Ma) because of unprecedented δ13C fluctuations to b − 10‰ PDB through thick (N1000 m) succession of stratigraphically complex sedimentary rocks deposited during tens of millions of years. In contrast, Phanerozoic large negative C-isotope excursions have been interpreted as the result of diagenetic fluid mixing during carbonate stabilization and burial and are less enigmatic due to the excellent biostratigraphic control on their timing and duration. The Ediacaran Nafun Group of Oman (part of the Huqf Supergroup spanning the Cryogenian–Early Cambrian) contains a large δ13C negative excursion (the Shuram excursion) reaching values as negative as − 12‰ at the base of the Shuram Formation. A steady recovery to positive values occurs over the entire Shuram and half through the overlying Buah Formation, suggesting a duration on the order of tens of My. Based on trace metal, chemostratigraphic and sedimentological analyses, the carbon isotope record obtained from the Buah Formation of northern Oman indicates a systematic and reproducible shift of δ13C values from − 6‰ to + 1‰ in 1 — a demonstrably diagenetic altered carbonate-cemented siliciclastic facies, and 2 — a least diagenetically altered stromatolitic facies. The identical reproducible isotopic pattern in these timeequivalent sections combined to the presence of exceptionally preserved δ18O values around − 2 to + 1‰ associated with the most negative δ13C values rules out isotopic resetting by diagenetic fluids as a mechanism to explain these values. It is concluded that it is possible to retain depositional δ13C values in demonstrably diagenetically altered carbonates. This raises the issue of the ability to recognize diagenetic alteration of C-isotopic values in Neoproterozoic rocks where a robust time frame to support reproducibility is not available. The results of this study provide strong support to a non diagenetic origin of the negative Shuram C-isotope excursion, believed to be the most profound (in terms of amplitude and duration) in the Earth's history. © 2009 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.

1. Introduction Chemostratigraphy serves as one of the principle means of intraand extra-basinal stratigraphic correlation used to assemble the fragmentary Precambrian stratigraphic record, compensating for the poor biostratigraphic resolution of Precambrian fossils (Knoll et al., 1986; Kaufman et al., 1997; Corsetti and Kaufman, 2003; Halverson et al., 2005; Maruyama and Santosh, 2008; Uchio et al., 2008; Le Guerroué, in press). For example, a negative carbon isotopic anomaly serves as one of the key criteria for defining the GSSP for the recently introduced Ediacaran Period of the Neoproterozoic Era (Knoll et al., 2006), because of its distinctive magnitude and occurrence at a precise stratigraphic position. The Shuram δ13C negative excursion (Halverson et al., 2005; Le Guerroué et al., 2006c) is the largest (in ⁎ Corresponding author. Beicip-Franlab. 232, Avenue Napoléon Bonaparte, PO Box 213, 92502 Rueil-Malmaison, France. Tel.: +33 672040353; fax: +33 1 47 08 41 85. E-mail address: [email protected] (E. Le Guerroué).

terms of amplitude and duration) of all Neoproterozoic 13C shifts and represents a fundamental challenge of the current understanding of the carbon cycle that should not be able to sustain δ13C values ≤5‰ for long periods of time (Des Marais and Moore, 1984; Melezhik et al., 2001; Le Guerroué and Kennedy, 2007). However, negative PDB δ13C (especially if coupled to negative δ18O PDB values; Gross and Tracey, 1965; Banner and Hanson, 1990) are also indicative of diagenetic processes (Kaufman et al., 1991). No known Precambrian succession shows subsequent positive δ18O associated with negative δ13C (see compilation in Shields and Veizer (2002)). Instead, they often show a strong cross correlation although there are no known primary oceanographic mechanisms linking those isotopic systems. Alternatively, meteoric early diagenesis of carbonate rocks involving stabilization of the carbonate mineralogy mixed with isotopically light meteoric fluids results in a strong covariation of both isotopic systems. Ultimately Neoproterozoic carbon isotope excursions could be the outcome of local basinal diagenetic processes and have neither time significance nor oceanographic importance.

1342-937X/$ – see front matter © 2009 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved. doi:10.1016/j.gr.2009.11.002

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In the Phanerozoic, the good time calibration of chemostratigraphic sections allows to indentify diachronous isotopic shifts representing diachronous diagenetic surfaces (Swart and Eberli, 2005). In the Precambrian, because of the lack of radiometric and biostratigraphic time constraints diagenetic surfaces can be less evident to recognize. Since common diagenetic processes drive δ13C and δ18O values towards more negative values, caution needs to be taken in addressing Neoproterozoic values as primary (Kaufman et al., 1993; Ishikawa et al., 2008; Melezhik et al., 2008a; Le Guerroué, in press). Critical to the veracity of these δ13C excursions are 1) identifying the origin of the carbonate material; 2) the reproducibility and synchronicity of these excursions; and 3) the δ13C dependence to sedimentary facies and isotopic shifts within homogenous sedimentary facies to clearly establish potential diagenetic issues linked to sedimentary facies. 2. Global carbon mass balance Carbon isotope ratios from marine carbonates precipitated in wellmixed surface water (above the chemocline) can record the global burial fraction of carbonate carbon and organic carbon, fixed by photosynthesis (Broecker and Peng, 1982). Photosynthetic activity in surface water preferentially removes 12C into organic matter, increasing 13C/12C that is then incorporated into carbonates (isotopically heavier values). Cessation of photosynthetic activity results in no preferential 12C removal from marine water into organic carbon and so carbonates assume a value of ∼−5‰, which is the unmodified bulk average value (Kroopnick, 1985). If photosynthesis ceases (the ‘Strangelove ocean’; Kump, 1991), then the global ocean value returns to the mantle value (∼−5‰) in a few hundred thousand years as marine carbon is replaced by riverine carbon (Kump, 1991; Bartley and Kah, 2004). Ultimately, the fractionation of carbon by photosynthesis accounts for the isotopic variability in sedimentary rocks and controls the steady-state values. Short term disruption (b106 years) of ecosystems is possible from events such as the effects of bolide impact (Olsen et al., 2002) or massive methane clathrate destabilizations (Dickens et al., 1995; Hesselbo et al., 2000). There are no well accepted mechanisms that can cause the complete cessation of primary productivity across the Earth for many millions of years. For carbonate values providing a (primary) record of the ocean/ atmosphere system, strong isotopic variability (non-steady-state) is limited by mass balance arguments to the residence time of carbon in the ocean, which is in the order of a few hundred thousand years. The most negative isotope excursion in Fig. 1C (e.g. the “Shuram excursion”) fundamentally challenges the current understanding of isotopic mass balance as established in the modern ocean (Broecker and Peng, 1982). The long time-scale required by the stratigraphic complexity of the Shuram excursion, largely longer than the residence time of carbon in the ocean (∼ 105 years), implies that these strongly negative isotopic values (below −10‰) record a sustainable steadystate condition in the global ocean (Le Guerroué et al., 2006c; Le Guerroué, in press). This contrasts sharply with the rare, short-lived, excursions of only 4‰ accepted in the Phanerozoic that have been used as evidence only for transient (short term) global ecosystem disruptions (Veizer et al., 1999; Hayes et al., 1999; Zachos et al., 2001) while large negative values recorded in Phanerozoic carbonates, have been interpreted as diagenetically altered (Fig. 2; see good example in Swart and Eberli (2005)). 3. Diagenetic patterns Almost all common diagenetic processes shift carbon and oxygen isotopes to more negative values as carbonate minerals stabilize in the presence of fluids that are 13C and 18O depleted (Fig. 2; Gross and Tracey, 1965; Hennessy and Knauth, 1985; Swart and Eberli, 2005).

Possible diagenetic processes include decarboxylation during deep burial of carbonate and organic matter, early organic diagenesis, or meteoric alteration during carbonate stabilization. This evidence includes 1) a strongly depleted oxygen isotope values that covary with δ13C values (Fig. 2; Swart and Eberli, 2005), 2) diagenetic textures including concretions, burial cements, dolomitization that present mixed isotopic values with the bulk rock samples (Figs. 2 and 3; Gross and Tracey, 1965; Irwin et al., 1977), 3) facies dependence, and an isotopic covariation with percent carbonate (Stephens and Sumner, 2003), 4) association of negative values below exposure surfaces with steps to heavy values above (Gross and Tracey, 1965; Swart and Eberli, 2005), and 5) strongly depleted oxygen isotope values indicating an open system for fluids as isotopic exchange happens through carbonate re-crystallization (Hennessy and Knauth, 1985; Kaufman et al., 1993; Melezhik et al., 2005). The carbonate mass balance is thought to be dominated by marine sources of carbonate (Kaufman et al., 1991) and the system is relatively closed to fluids with sufficient carbon to alter the isotopic ratios (this is based on the modeling of Banner and Hanson (1990)). From Phanerozoic studies, however, thick intervals of platformal carbonates can be altered during CaCO3 stabilization. These studies show that the original transformation of primary precipitate to limestone/dolostone in shallow-water environments commonly occurs in the presence of meteoric waters (Fig. 2). Low 18O meteoric fluids enter the system from land areas along coastlines, bathe subaerially exposed deposits, and move into offshore deposits by groundwater flow and/or coastal progradation. Meteoric waters descending through land surfaces can become strongly depleted in 13 C relative to marine pore fluids from microbial respiration of photosynthetically fixed carbon. Carbonates stabilized in such fluids show marked depletion in 13C and 18O with strong covariation in δ18O and δ13C and are well understood from a large number of studies of Phanerozoic carbonates like the Unda and Clino sections of the Great Bahamas Bank (Swart and Eberli, 2005, Fig. 2) or studies on Miocene carbonates (Hennessy and Knauth, 1985). The open system nature of diagenesis is particularly apparent where thick platformal successions are dolomitized and Ca is replaced by Mg in low Mg concentrated pore fluids (the so-called ‘dolomite problem’, e.g. Morrow, 1990). Phanerozoic example presents, in vertical section, highly variable non consistent values (attaining − 8‰; Fig. 2) constrained below sequence boundaries. 4. Geological background 4.1. The Shuram and Buah formations sedimentology The sedimentology and isotope record of the Shuram and Buah formations have been extensively described in many studies (Burns and Matter, 1993; McCarron, 2000; Cozzi and Al-Siyabi, 2004; Cozzi et al., 2004a,b; Le Guerroué et al., 2006a,b,c and references therein). The reader is referred to these articles for more details. Key observations relevant to the discussion and new evidences are presented in this work. The Nafun Group of Oman is made of two clastic–carbonate sedimentary cycles which span the whole Ediacaran Period. The second of these cycles is made of the Shuram (clastic) and Buah (carbonate) formations, for a total thickness of 600–1000 m. Outcrops in the Jabal Akhdar mountains (northern Oman) show prevailing deeper water facies with respect to the Huqf area ones (central Oman) where shallow-water facies prevail. The Shuram Formation is a thick shelfal sandy/silty succession with an upward increase in carbonate content. It is a compelling example of an operating mixed carbonate/ siliciclastic system, with deposition occurring on a storm-dominated, well-mixed shelf. Mudstones, ooids and sand grains are interbedded and mixed at every scale and deposited in high-energy cross-bedded packages (Le Guerroué et al., 2006b). The carbonate-rich units are

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Fig. 1. End of the Shuram δ13C excursion in northern Oman (A and B) recorded in the Buah Formation (C) in the Wadi Hajir (D) and Wadi Bani Awf (E) sections. A: Localization map showing Neoproterozoic outcrops. B: Simplified geological map showing the studied sections. C: Shuram δ13C excursion (Le Guerroué et al., 2006c). Dashed box represents studied stratigraphic interval. D and E: Detailed δ13C stratigraphic record presenting a sharp shift in δ13C values in very different sedimentary facies. Isotope data are after Cozzi et al. (2004a) and Cozzi et al., 2004b (circles) and from this study (lozenges).

found in the upper part of asymmetric, m-scale siliciclastic/carbonate cycles with increasing carbonate towards the top. Massive m-thick storm-dominated siliciclastics pass to cemented discontinuous cmthick beds of siltstones often capped by cm-thick oolite beds. The

Shuram Formation gradually passes into the limestone/dolostone Buah Formation via a 20–30 m-thick transition zone where stormgenerated carbonate beds and red siltstone lithologies alternate, with the former becoming more abundant towards the base of the Buah

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Fig. 2. Standard δ18O/δ13C cross plot for demonstrated Phanerozoic diagenetic textures including: Cemenstones (R2 = 0.78, n = 29; Irwin et al., 1977), the Rannoch Formation stratabound (R2 = 0.1475, n = 38; Prosser et al., 1993), meteorically altered carbonate ramp from the Unda (R2 = 0.6109, n = 437) and Clino (R2 = 0.7419, n = 637) sections of the Great Bahamas Bank (Swart and Eberli, 2005) and the Bikini and Eniwetok Atolls (R2 = 0.5948, n = 36; Gross and Tracey, 1965). Note the strong mixing relationship (high R2 values) between primary marine values (positive carbon and oxygen isotopes) and meteoric values (very negative carbon and oxygen isotopes).

(McCarron, 2000; Cozzi and Al-Siyabi, 2004; Cozzi et al., 2004a,b; Le Guerroué et al., 2006b). The Buah Formation was deposited on a storm-dominated distallysteepened carbonate ramp. In the Jabal Akhdar outcrops of northern Oman mid-outer ramp facies are found in Wadi Bani Awf and inner ramp facies are present in Wadi Hajir and in the south-eastern sections (Cozzi et al., 2004b). Optimal carbonate production zone in the Buah carbonate ramp depositional system was located in the midinner ramp where stromatolites and ooids accumulated as a thick succession (Fig. 3C). In Wadi Bani Awf, above condensed bituminous mudstones, resedimentation processes dominated, with abundant slumping and brecciation of outer ramp sediments and redeposition on a carbonate slope, as well as sediment input from the inner ramp facies (rudstone–floatstone bed; Cozzi et al., 2004b). This important sedimentary facies and stratigraphic evolution difference is due to an initially inclined depositional surface gently sloping towards Wadi Bani Awf at the end of Shuram time as deduced from the presence of slightly shallower water facies in Wadi Hajir than in Wadi Bani Awf (Cozzi et al., 2004b). This differentiation into a deeper setting in Wadi Bani Awf and shallow-water setting in Wadi Hajir (generally east and south) continued until the end of Buah deposition and was probably accentuated by synsedimentary tectonism which affected the Buah Formation starting from its middle part, determining an increase in subsidence rates to the west and north. Despite this fact, the upper part of the Buah in the Wadi Bani Awf and Wadi Hajir sections both record a shallowing upward trend (Cozzi et al., 2004b) which is believed to be the result of deposition during a Highstand Systems Tract (HST), in accordance with the vertical facies stacking patterns of outcrops and subsurface sections throughout the Oman Ediacaran Basin (Cozzi et al., 2004a). 4.2. The Shuram and Buah formations isotopic record The general pattern of variation in δ13C of the Shuram excursion is largely reproducible throughout Oman, including outcrop and

Fig. 3. Field pictures of the Buah Formation facies. A) Carbonate-cemented siliciclastics interbedded with carbonate bands of the Wadi Bani Awf section. B) Carbonate concretionary lens details found in A. C) Columnar stromatolitic facies found in Wadi Hajir. See further detail in text. Lens cap is 50 mm in diameter and coin is about 15 mm.

subsurface (exploratory wells) data, and irrespective of sedimentary facies (Fig. 1C; Burns and Matter, 1993; McCarron, 2000; Cozzi and AlSiyabi, 2004; Fike et al., 2006; Le Guerroué et al., 2006a,c; Le Guerroué, in press). Isotopic values start positive in the underlaying Khufai Formation slowly reaching zero towards its top. δ13C values shift progressively from 0 to −8‰ over 3 m of peloidal carbonate facies (Le Guerroué et al., 2006a). In the siliciclastic-dominated Shuram Formation δ13C values drop further to reach −12‰. The δ13C record then steadily moves towards less negative values eventually passing into the Buah Formation and returning to positive values in the upper half of this formation (Fig. 1). 5. Material and methods To investigate potential diagenetic alteration issues of the Shuram excursion two stratigraphic and time equivalent, homogeneous but profoundly different sedimentary facies containing a large variation in δ13C were targeted in this study. The facies homogeneity is important because one would assume diagenetic processes to be dependent on facies characteristics. For example, the C-isotope shift could be linked to a porosity gradient throughout the facies. This sedimentological data is complemented by new chemostratigraphic and trace metal data, obtained in the course of measuring stratigraphic sections. If the

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large δ13C variation is the result of diagenetic alteration then the homogenous sedimentary facies will allow this hypothesis to be tested, and a gradient in the diagenetic overprint should appear and correlate to more negative δ13C values. Geochemical samples were drilled with 1–2 mm dental drill bits from freshly cut rock slabs avoiding sparry cement and vein material. The C- and O-isotope composition of powder from the carbonate samples was measured with a GasBench II connected to a Finnigan MAT DeltaPlus XL mass spectrometer, using a He-carrier gas system according to the methods adapted after Spoetl and Vennemann (2003). Isotopic compositions are reported in the δ-notation relative to VPDB (Vienna PeeDee belemnites) for carbon and oxygen. Samples are normalized using an in-house standard calibrated against δ13C and δ18O values of NBS-19 (+1.95 and −2.20‰, relative to VPDB). External reproducibility for the analyses estimated from replicate analyses of the in-house standard (n = 6) is ±0.07‰ for δ13C and ±0.08‰ for δ18O. Results are shown in Table 1. Powders from the same drilled samples were used for major and trace elements of the carbonate phase. Samples were leached in 0.2 N HCl for 2 h at room temperature, centrifuged and then analyzed at the University of California at Riverside (UCR) using a Perkin Elmer 3000DV inductively coupled plasma optical emission spectrograph. The precision (1σ) is typically around 2% of the major oxide present. Precision of the measurements is about 2% for Ca and 5% for Mg and Sr. All measured concentrations were well above the lower limit of detection. Results are shown in Table 1.

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numerous (∼ 60) exploratory wells that encountered the Buah Formation (Cozzi and Al-Siyabi, 2004) and 3) an identical, reproducible throughout the Oman Ediacaran Basin, shift in C-isotopes from −6‰ to + 1‰ recorded in homogenous facies (Figs. 1 and 3; Cozzi et al., 2004a,b). This isotope shift corresponds to the end of the large Shuram excursion (Le Guerroué et al., 2006c). 6.1. Wadi Bani Awf section The first section located in Wadi Bani Awf presents in its middle (around 60 m from the base of the formation) m-thick cycles of carbonate-cemented thin siltstones interbedded with cm-thick disrupted bands of dolostones (this is the upper part of the 20 m-thick rudstone–floatstone succession in Cozzi et al. (2004b)). While some carbonate bands are concretionary in origin, most of them are detrital in origin and have been interpreted as distal, super-fine calcareous turbidites (Cozzi et al., 2004b), which were subsequently deformed by boudinage processes (Fig. 3A and B). The carbonate material is made of stoichiometric re-crystallized dolomicrites. Thin section petrographic analyses failed to reveal any recognizable carbonate allochems, and the origin of the primary carbonate grains remains unknown. The last of these cycles (4 m-thick) bares an important δ13C variation from about − 6‰ to + 1‰ (Fig. 1 and Table 1). Cemented siltstones, concretions and reworked floatstones share similar isotopic values with respect to their stratigraphic position (Table 1). Concretions failed to demonstrate any δ13C gradient from core-to-rim that is a general observation in the Shuram and Buah formations (Table 1; Le Guerroué and Kennedy, 2007). Trace metal composition reveals quite some covariation between elements such as Mn, Fe and Mn/Sr against δ13C. This facies also has the highest Mn/Sr ratio (Table 1 and Fig. 4). Mg/Ca does not appear to covary with δ13C. Carbonate abundance in bulk sample covaries against δ13C for this very specific facies (Fig. 4) but elsewhere in the Shuram or Buah formations it does not (data in Le Guerroué et al. (2006a)). δ13C/δ18O cross plot (Fig. 5) shows quite some covariation that is a general trend in the Buah Formation in the area (R2 = 0.72) although the trends are different. However, this facies presents unusually high δ18O values (between −2 and 0‰) associated with some (but not all) of the most depleted 13C samples (Fig. 5).

6. Sedimentological and chemostratigraphic results Two specific stratigraphic intervals in the Buah Formation located in the Jabal Akhdar of northern Oman were investigated. Both sections are considered to be time equivalent based on three main datasets: 1) the identical stratigraphic position, the Buah Formation being sandwiched between the reddish siltstones of the Shuram Formation at the base and the Fara Fm volcaniclastics at the top (Cozzi et al., 2004b); 2) the identical vertical stacking pattern of sedimentary facies which reflects deposition during HST conditions (see Section 4.1 for a more detailed explanation) identified not only in the Jabal Akhdar sections but also in the Huqf outcrops of central Oman as well as in the

Table 1 Geochemical composition. Key is WBA: Wadi Bani Awf; WH: Wadi Hajir. Stratigr. Height (m)

Sample name

Locality

0.0 0.4 0.8 0.8 0.8 1.2 1.6 2.0 2.4 2.8 3.2 3.6 4.0 4.4 0.0 2.0 4.0 5.9 7.3 7.9 9.2 9.5 10.1 12.0

Bsil-1 Bsil-2 Bsil-3a Bsil-3b Bsil-3c Bsil-4 Bsil-5 Bsil-6 Bsil-7 Bsil-8 Bsil-9 Bsil-10 Bsil-11 Bsil-12 Bstr-1 Bstr-2 Bstr-3 Bstr-4 Bstr-5 Bstr-6 Bstr-7 Bstr-8 Bstr-9 Bstr-10

WBA WBA WBA WBA WBA WBA WBA WBA WBA WBA WBA WBA WBA WBA WH WH WH WH WH WH WH WH WH WH

Lithology

Limestones Cemented Silst. Concretion Concretion Concretion Limestones Cemented Silst. Limestones Limestones Limestones Cemented Silst. Limestones Limestones Limestones Stromatolites Stromatolites Stromatolites Stromatolites Stromatolites Stromatolites Stromatolites Stromatolites Stromatolites Stromatolites

Wt.% MgCO3

% Carbo

Isotope δ C

δ O

Ba

Ca

Fe

K

Mg

Mn

Sr

Zn

49% 46% 47% 47% 47% 46% 48% 48% 47% 48% 47% 47% 48% 49% 49% 50% 50% 48% 49% 49% 49%

100 39 68 65 67 81 34 97 64 81 44 67 83 94 100 100 100 100 100 100 100 100 100 100

− 6.1 − 6.1 − 6.5 − 6.4 − 6.5 − 6.0 − 5.8 − 5.7 − 5.1 − 4.1 − 6.0 − 3.6 − 3.4 − 3.3 − 5.3 − 5.6 − 5.3 − 4.9 − 0.8 − 1.0 0.6 0.2 − 0.2 − 2.3

− 3.4 − 1.1 − 2.2 − 1.8 − 2.2 − 2.0 − 0.8 − 1.4 − 1.5 − 1.2 − 0.5 − 1.4 − 1.9 − 1.7 − 7.1 − 3.0 − 3.4 − 4.0 0.5 − 0.2 0.1 0.4 − 1.1 − 0.1

0.45 4.34 4.98

4.67 4.83 4.83

4.68 16.07 14.66

0.18 0.84 0.40

5.26 4.94 4.99

2.21 6.76 3.44

0.57 1.19 1.19

0.45 1.32 0.77

8.59 2.21 0.65 1.67 1.08 4.96 1.84 0.85 0.84 1.56 0.09 0.89 2.34 0.00 0.22 0.00

4.94 4.69 4.68 4.83 4.77 4.81 4.85 4.73 4.67 4.66 4.59 4.62 4.77 4.63 4.65 4.67

8.06 10.18 8.92 5.63 4.41 12.04 3.06 3.00 2.63 0.32 0.15 0.20 0.19 0.40 0.50 0.85

0.68 0.28 0.34 0.33 0.38 0.47 0.45 0.24 0.32 0.14 0.05 0.03 0.06 0.11 0.13 0.16

4.95 5.18 5.20 5.10 5.16 5.03 5.10 5.22 5.29 5.33 5.41 5.38 5.23 5.36 5.34 5.32

2.07 2.78 2.45 1.63 1.80 4.17 1.60 1.35 1.16 2.26 1.19 0.49 0.90 0.00 0.00 0.04

1.26 0.87 0.89 1.08 0.90 1.34 1.07 0.85 0.79 0.41 0.30 0.55 0.50 0.78 0.60 0.58

6.48 5.03 0.42 1.58 5.55 1.00 2.14 2.92 3.02 1.10 0.41 0.50 0.63 0.18 0.25 0.74

0.84 1.19

4.59 4.61

0.67 0.51

0.11 0.15

5.40 5.39

2.30 0.17

0.56 0.73

0.48 0.30

50% 50%

13

Carbonate phase relative mol.% 18

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Fig. 4. Geochemical data of the Buah Formation in facies presenting the sharp shift in their δ13C record (Fig. 1). A: C–O-isotopes (VPDB), carbonate weight (carb.), MgCO3 relative weight and relative mole abundance of different elements (normalized to carbonate content) represented against stratigraphic height. B: same data is plotted against δ13C. Grey arrows represent diagenetic trends. Note the very distinctive diagenetic signature between the two facies.

6.2. Wadi Hajir section The second section crops out in Wadi Hajir and occurs at around 60 m from the base of the Buah Formation massive 60 m-thick dolomitized succession of meter-scale stromatolite mounds, with both elongated and more symmetrical morphologies corresponding to mid and inner carbonate ramp facies (Figs. 1 and 3; Cozzi et al., 2004b). There is no siliciclastic input in the stromatolitic facies that is also completely dolomitized (50% MgCO3). Higher up the dolomitization is more partial (Fig. 4; Cozzi et al., 2004b). Towards its top, this facies records the same shift in δ13C as the one in the Wadi Bani Awf (values rising from −6‰ to +1‰ VPDB within 9 m of stratigraphy; Figs. 1 and 4). Trace metal content does not show

enrichment in Fe, Zn or Mn that does not covary with the δ13C values. The Mn/Sr ratio is lower than in the other facies, well below 3. δ13C/ δ18O cross plot (Fig. 5) shows quite some covariation but this is not a trend at the section scale (R2 = 0.52). However, this facies, like the one in Wadi Bani Awf, presents high δ18O values (between −2 and +1‰) associated with some of the most depleted 13C samples (Fig. 5). However, some of the most depleted 13C values do fall on trend of the general δ13C/δ18O covariation line. This facies contrasts sharply with the one in Wadi Bani Awf in terms of origin of its carbonate material and trace metal composition. Still they both display a very similar δ13C pattern and in both facies, the stratigraphic interval showing the rapid shift in δ13C (from − 6‰ to +1‰; Fig. 1) is the one recording the least depleted 18O values (Fig. 5). These values are unique to the Shuram excursion data set (McCarron, 2000; Fike et al., 2006; Le Guerroué et al., 2006a,c). 7. Discussion 7.1. Source of the carbonate

Fig. 5. Standard δ13C/δ18O cross plot for the Buah Formation including WBA (Wadi Bani Awf) and WH (Wadi Hajir) studied facies. Note the overall quite good covariation (R2) and the presence of δ18O values in the − 2‰ to + 1‰ area associated with very negative δ13C values (around − 6‰) contrasting sharply with Fig. 2.

Chemostratigraphy assumes that the δ13C values are obtained from carbonate precipitated in the well-mixed, isotopically homogenous portion of the ocean above the chemocline (Brand and Veizer, 1981). This is a particular problem in the Precambrian because the absence of skeletal carbonate grains precipitated in a known environment and water depth. Carbonate lithologies, such as micrite and detrital carbonate silt from which the bulk of Precambrian values are derived, show greater variability when similar lithologies are compared to cooccurring pelagic skeletal precipitates in Phanerozoic sediments (Stephens and Sumner, 2002). This is an even greater problem with a Precambrian succession such as the Wonoka (Australia) and Shuram

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formations where values are recorded within a specific facies of thick, mixed siliciclastic–detrital carbonate (Le Guerroué and Kennedy, 2007). A primary origin of the carbonate material, well within the homogenous photic ocean is obvious for the case of the shallow-water stromatolitic mounds of the Wadi Hajir section. However, in the siliciclastic-dominated cycles of the Wadi Bani Awf section a primary source of carbonate material can be argued for. Carbonate originating from the shallower parts of the carbonate ramp was reworked by storm activity and re-deposited as debris sheets in the outer ramp setting, attesting for an early cementation prior to reworking, producing carbonate clasts (Cozzi et al., 2004b). Alternatively, later carbonate cementation of the siliciclastic horizons in Wadi Bani Awf is attested by the presence of numerous concretions that do not show concentric growth and could be the result of pervasive diagenetic fluids bathing the Buah carbonate ramp like other Phanerozoic examples (Figs. 2 and 3). Eventually, isotopic similarities between concretions, cemented siliciclastics and storm reworked material attest for a similar origin (or later pervasive alteration) for most of the carbonate material (Fig. 4 and Table 1). If a primary carbonate source hypothesis is considered, then the carbonate cementation has to be originated from the dissolution/re-crystallization of primary carbonate components. The presence of δ18O values close to zero (refuting diagenetic alteration; Fig. 2) throughout the Wadi Bani Awf and Wadi Hajir facies supports this primary depositional origin (Fig. 5) and argues against diagenetic resetting (Fig. 2). Although later diagenetic alteration is obvious in the Wadi Bani Awf section (see discussion below), the primary carbonate was able to retain depositional isotopic composition and largely buffered later diagenetic fluids. 7.2. Diagenetic alteration Mn and Sr abundances and oxygen isotope compositions of samples are of particular interest as these can be redistributed under the influence of meteoric fluids alteration (Brand and Veizer, 1980; Veizer, 1983). The Mn/Sr ratio is a common proxy for diagenetic exchange. During meteoric diagenesis Sr is expelled from marine carbonate, whereas Mn is incorporated (Derry et al., 1992; Kaufman and Knoll, 1995). The Mn/Sr ratio in Buah dolomicrites shows values well under 3 in the Wadi Hajir section, whereas the Wadi Bani Awf one shows ratios between 3 and 10 (Fig. 4). This is in agreement with more general study on the Shuram excursion that presents most samples with Mn/Sr ratios b10, and many samples b3 (Fike et al., 2006, Le Guerroué et al., 2006a). Carbon isotopes from samples in Wadi Hajir may be considered as less altered than the ones in Wadi Bani Awf given the lack of relationship between δ13C and Mn/Sr values (Fig. 4). A diagenetic trend in the Wadi Bani Awf section is also supported by quite some correlation between elements such as Fe, Zn and Mn and the C-isotope ratio whereas the Wadi Hajir samples lack covariation (Fig. 4). When the plot of δ13C versus δ18O (Fig. 5) produces a straight line of positive slope, the covariance is thought to be due to meteoric alteration, which reduces both the carbon and oxygen isotopic ratios (Gross and Tracey, 1965; Fairchild et al., 1990; James and Choquette, 1990; Marshall, 1992; Figs. 2 and 5). Commonly in the carbonate system, δ18O signal is likely to be altered due to high water-to-rock ratios with host-rock not capable of buffering fluid δ18O values (Kaufman et al., 1991; Corsetti and Kaufman, 2003). The δ18O signature of dolomites especially is generally influenced by one of the fluids because dolomitization requires large volumes of water to supply adequate Mg for the newly formed mineral (Banner and Hanson, 1990). As a result, there is a paucity of values retaining original δ18O values in diagenetic systems (Gross and Tracey, 1965; Brand and Veizer, 1981; Swart and Eberli, 2005). This is particularly the case in Neoproterozoic carbonates (see compilation in Veizer et al. (1999), Shields and Veizer (2002) and reference therein).

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In the overall Buah Formation the most negative δ13C values are associated with the most negative δ18O ratio suggesting variable degrees of alteration (Fig. 5; R2 = 0.52 and 0.72). However, the smooth and steady rise in δ13C values in closely spaced samples argues against a diagenetic alteration that predicts a high variable record below a sequence boundary (see good example in Swart and Eberli (2005)). The O-isotope record also shows values very close to zero attesting to a minimal diagenetic alteration of the isotopic system (Fig. 5). The association of these unique, close to zero δ18O values with δ13C values around − 6‰ (this is not the case of all values) suggests that the C-isotope ratio must also retain a primary depositional signal since the buffering effect of carbonate carbon relative to the bicarbonate abundance of diagenetic solutions is stronger (Kaufman et al., 1991; Corsetti and Kaufman, 2003). However, it must be noticed that generally the most negative δ13C values (−12 to −8‰) in the lower Buah and the Shuram formations do not present such high δ18O values (Fig. 5). This remarkably preserved δ18O values validate the presence of primary depositional −6‰ δ13C values at the end of the Shuram excursion and provide great support for the overall anomaly to be the record of a primary oceanographic signal.

7.3. The veracity of the Shuram excursion Interpreting the b − 10‰ δ13C Shuram excursion (Fig. 1) as primary versus diagenetic in origin is not straightforward. A primary origin is supported by the good reproducibility of the excursion in different basins and the presence of second order shifts occurring simultaneously in different homogenous facies (Figs. 1 and 4; see discussion above). The diagenetic origin is supported by the very negative nature of the isotopic values and a good cross correlation between δ13C and δ18O (Fig. 5) especially in the carbonate-cemented siliciclastic facies that makes the most of the Shuram Formation (Le Guerroué and Kennedy, 2007). At the formation scale, it appears that lithologies coincide with isotopic values (Fig. 1C). This is a common problem to the Wonoka and Shuram formations that present their most negative δ13C values within siliciclastic-dominated facies (see comprehensive studies in Calver (2000) and Le Guerroué et al. (2006a,c)). Where the source of the carbonate is obvious in the form of stromatolites and oolites, then the carbonate δ13C values tend to be positive (but see the −9‰ δ13C oolitic shallow-water carbonates of the upper Shuram Formation; Le Guerroué et al., 2006b). Very negative values appear within the first cemented siltstones of the Shuram and return to positive values at the top of the Buah Formation (Fig. 1). A similar situation appears in the lower Wonoka Formation that returns to positive values (+1 to +2‰) in the upper unit no. 11 where stromatolites and oolites are present, pointing towards diagenetic alteration of the cemented siliciclastics lower parts (Le Guerroué and Kennedy, 2007). However, there is no clear relationship between carbonate abundance in bulk samples and 13C depletion, and at the scale of the global Shuram excursion, intervals defined by strong negative δ13C values contain similar carbonate abundances to those reflecting no isotope depletion (see compilation in Le Guerroué et al. (2006a)). Since diagenetic alteration produces recognizable isotopic patterns (see above) and is intimately linked to meteoric fluid bathing, the test is to look for repeated isotope variations around individual carbonate cycles. Le Guerroué et al. (2006b) demonstrate that a few eustatic cycles (shallow-water parasequences presenting δ13C values around −9‰) at top Shuram support a coherent secular δ13C variation spatially distributed along individual parasequence progradation direction and vertically within the stack. This result rules out facies/ porosity controlled diagenetic processes such as meteoritic fluid mixing during eustatic cycles and validates − 9‰ δ13C values as primary.

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Eventually, critical to the veracity of the Shuram excursion as a primary carbon cycle perturbation is its global reproducibility and synchronicity. Many sections in Oman, Siberia, Australia, China and further more do show similar δ13C values in their Ediacaran succession suggesting tentative correlations. However, the absence of a robust biostratigraphic calibration and the general lack of radiometric age constraints do not allow the synchronicity issue to be addressed with precision. The use of the strontium isotope ratio as an independent chronometer provides a good alternative and suggests a possible global δ13C perturbation starting around 590 Ma and ending around 550 Ma (Le Guerroué, in press). Many recent studies acknowledge a probable global excursion (Zhu et al., 2007; Melezhik et al., 2008b) but disagreement persists on the age of the onset of the excursion. Propositions are 600 Ma based on basin subsidence analyses and detrital zircons analyses (Le Guerroué et al., 2006c) or younger ages around 580–555 Ma based on correlation with the Gaskiers glaciation or the end of the δ13C excursion constrained in China (Halverson et al., 2005; Condon et al., 2005; Bowring et al., 2007). A whole ocean isotope excursion hypothesis faces three problems: 1) the Shuram excursion values are far more negative than the − 5‰ oceanic value that defines an absolute lower limit in the Phanerozoic (Shields and Veizer, 2002); 2) they are sustainable through thick packages of sediments (Fig. 1) most likely deposited over many millions of years (Halverson et al., 2005; Zhou and Xiao, 2007) and probably up to a few tens of millions of years (Le Guerroué et al., 2006c; Le Guerroué, in press) implying steady-state conditions. 3) Until today there is no mechanism capable of explaining the observed correlation between most negative δ13C and δ18O signals at the global scale. If the Shuram excursion is secular and representative of the changing global ocean value in steady-state conditions then the exogamic cycle must have functioned with very different, and currently unrecognized, controls than during the Phanerozoic involving carbon reservoir with much longer residence time (Rothman et al., 2003). The size and consistent supply of the 12C pool capable of shifting the isotopic value of the ocean to b − 10‰ over this period of time is problematic. If rapid oxidation of a larger pool of dissolved organic carbon hypothesized for the Precambrian oceans provides a compelling explanation for the potential to achieve very negative oceanic values for short time intervals (Rothman et al., 2003; Fike et al., 2006), this process is a non-steady-state reallocation of carbon within preexisting pools and is not likely sustainable over time scales of N107 years and is also limited by availability of oxidants. The primary interpretation of the end of the Shuram excursion suggested by this study has a significant implication for the interrelationship of evolutionary events and the carbon cycle that has not been explored (Fike et al., 2006; Peltier et al., 2007). To understand the excursion will require modeling carbonate precipitation in absence of organisms capable of secreting shells and could involve different behavior/reactivity of the oceanic dissolved organic carbon pool and probable longer residence time of methane in an atmosphere that probably was limited in O2 in the Neoproterozoic (Fike et al., 2006; Peltier et al., 2007). 8. Conclusions A critical test of the veracity of the very negative δ13C values making the Shuram excursion in Oman is the shifts occurring within homogenous sedimentary facies, because they would acquire a uniform diagenetic isotope ratio resetting. The two sections in the Buah Formation studied in this article, corresponding to the end of the Shuram excursion, represent a compelling example of important isotopic variation (N7‰) within 1 — a demonstrable concretionary diagenetically altered facies (Wadi Bani Awf section) and 2 — a massive stromatolitic facies (Wadi Hajir section). δ13C stratigraphic reproducibility combined to exceptionally high δ18O values (between −2 and + 1‰) preserved in these two facies attest to the preservation

of primary depositional isotopic values. This study demonstrates possible rock-buffering of original O- and C-isotope systems in diagenetically altered facies and provides support to the veracity of the Shuram excursion (δ13C b −10‰) as an oceanographic primary signal. However, this raises the question of why, at the large Shuram excursion scale, the δ18O signal somewhat cross correlates with the δ13C record (Fig. 5)? This observation remains to be constrained since it significantly resembles other diagenetically affected Phanerozoic carbonate ramps (Fig. 2). Interpreting negative isotope values as primary marine signals and correlating them is particularly problematic in Precambrian sediments. Much of what is sampled is not a recognizable allochem such as a shell or biogenic fragment with a clear marine origin. There is no independent basis for correlation of isotope values. Negative values caused by diagenetic alteration are more evident in Phanerozoic strata because biostratigraphy provides a robust time frame that allows to identify diachronous, diagenetically-made isotopic excursions. Negative isotopic values can occur quasi systematically below erosion and flooding surfaces where meteoric alteration is more common. The Wadi Bani Awf section demonstrates that preservation of primary isotopic values in diagenetically altered zones is possible. However, there is no satisfactory mechanism to explain the Shuram extraordinary negative excursion and higher standard of supports are necessary for confident interpretation of primary marine negative isotope values in the Neoproterozoic. Acknowledgements Swiss national research project (Number: PBEZ2-115170) is thanked for the financial support. We are especially grateful to Peter Swart for the use of his data in Fig. 2 and to Dave Mrofka for the endless help in the lab. This is a contribution to the IGCP-512 project. References Banner, J.L., Hanson, G.N., 1990. Calculation of simultaneous isotopic and trace element variations during water–rock interaction with applications to carbonate diagenesis. Geochimica et Cosmochimica Acta 54, 3123–3137. Bartley, J.K., Kah, L.C., 2004. Marine carbon reservoir, Corg–Ccarb coupling, and the evolution of the Proterozoic carbon cycle. Geology 32, 129–132. Bowring, S.A., Grotzinger, J.P., Condon, D.J., Ramezani, R., Newall, M., Allen, P., 2007. Geochronologic constraints on the chronostratigraphic framework of the Neoproterozoic Huqf Supergroup, Sultanate of Oman. American Journal of Science 307 (10), 1097–1145. Brand, U., Veizer, J., 1980. Chemical diagenesis of a multicomponent carbonate system; 1, Trace elements. Journal of Sedimentary Research 50 (4), 1219–1236. Brand, U., Veizer, J., 1981. Chemical diagenesis of a multicomponent carbonate system; 2, Stable isotopes. Journal of Sedimentary Research 51 (3), 987–997. Broecker, W.S., Peng, T.H., 1982. Tracers in the Sea. Eldigio Press, Palisades, New York. 690. Burns, S.J., Matter, A., 1993. Carbon isotopic record of the latest Proterozoic from Oman. Eclogae Geologicae Helvetiae 86 (2), 595–607. Calver, C.R., 2000. Isotope stratigraphy of the Ediacaran (Neoproterozoic III) of the Adelaide Rift Complex, Australia, and the overprint of water column stratification. Precambrian Research 100, 121–150. Condon, D., Zhu, M., Bowring, S., Wang, W., Yang, A., Jin, Y., 2005. U–Pb ages from the Neoproterozoic Doushantuo Formation, China. Science 308, 95–98. Corsetti, F.A., Kaufman, A.J., 2003. Stratigraphic investigations of carbon isotope anomalies and Neoproterozoic ice ages in Death Valley, California. Geological Society of America Bulletin 115 (8), 916–932. Cozzi, A., Al-Siyabi, H.A., 2004. Sedimentology and play potential of the late Neoproterozoic Buah carbonates of Oman. GeoArabia 9 (4), 11–36. Cozzi, A., Allen, P.A., Grotzinger, J.P., 2004a. Understanding carbonate ramp dynamics using δ13C profiles: examples from the Neoproterozoic Buah Formation of Oman. Terra Nova 16 (2), 62–67. Cozzi, A., Grotzinger, J.P., Allen, P.A., 2004b. Evolution of a terminal Neoproterozoic carbonate ramp system (Buah Formation, Sultanate of Oman): effects of basement paleotopography. Geological Society of America Bulletin 116 (11/12), 1121–1131. Derry, L.A., Kaufman, A.J., Jacobsen, S.B., 1992. Sedimentary cycling and environmental change in the Late Proterozoic: evidence from stable and radiogenic isotopes. Geochimica et Cosmochimica Acta 56, 1317–1329. Des Marais, D.J., Moore, J.G., 1984. Carbon and its isotopes in mid-oceanic basaltic glasses. Earth and Planetary Science Letters 69, 43–57. Dickens, G.R., O'Neil, J.R., Rea, D.K., Owen, R.M., 1995. Dissociation of oceanic methane hydrate as a cause of the carbon isotope excursion at the end of the Paleocene. Paleoceanography 10, 965–971.

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