Cambrian boundary in the South Oman Salt Basin, Sultanate of Oman

Marine and Petroleum Geology 39 (2013) 187e197

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Primary silica precipitate at the Precambrian/Cambrian boundary in the South Oman Salt Basin, Sultanate of Oman Karl Ramseyer a, *, Joachim E. Amthor b, Albert Matter a, Thomas Pettke a, Martin Wille c,1, Anthony E. Fallick d a

Institute of Geological Sciences, University of Bern, Baltzerstrasse 1þ3, CH-3012 Bern, Switzerland Petroleum Development Oman, PO Box 81, PC 113 Muscat, Oman Research School of Earth Sciences, Australian National University, ACT 0200, Australia d SUERC, Rankine Avenue, Scottish Enterprise Technology Park, East Kilbride G75 0QF, Scotland, UK b c

a r t i c l e i n f o

a b s t r a c t

Article history: Received 17 February 2012 Received in revised form 13 August 2012 Accepted 24 August 2012 Available online 5 September 2012

An organic-rich laminated chert (silicilyte) consisting of up to 90% microcrystalline quartz that formed at the PrecambrianeCambrian boundary acts as a light-oil reservoir in the subsurface of the South Oman Salt Basin, Sultanate of Oman. Fully encased in salt domes, it was first discovered during the 1990’s hydrocarbon exploration activities of Petroleum Development Oman. Because of the economic significance and the unconventional reservoir characteristics, there is great interest in understanding the origin of the laminated chert, its source of silica and mode of precipitation in an anoxic, sulphur-rich, stagnant and highly saline basin. The homogeneous distribution and high values of stable Si isotope composition (avg. d30Si ¼ 0.83  0.28) coupled with a low molar Ge/Si ratio (<0.25  106) of the microcrystalline matrix quartz clearly reveal dissolved silica in the seawater as the Si source, whereas hydrothermal or biogenic (e.g. sponge-derived) silica can be excluded. Silica precipitation from seawater was likely the result of a dramatic increase in salinity in response to salt dissolution atop and adjacent to the edges of transtensional depressions on the deep basin floor, thus markedly reducing the solubility of amorphous silica in these brine-filled seawater depressions. This saturation triggered the formation of silica-gel, which accumulated at the basin floor forming a soft silica-rich layer on bacterial mats giving rise to a laminated sediment. The mean number of laminae is ca. 32 per year suggesting that layering is non-annual and controlled by processes such as fluctuations in nutrient supply, lunar driven re-mixing or diagenetic segregation. The transformation of the silica-gel to microcrystalline quartz occurred below 45  C indicating a less than 4.5& d18O composition of the pore-water during microcrystalline quartz formation. Ó 2012 Elsevier Ltd. All rights reserved.

Keywords: Silicilyte Chert PCeC boundary d30Si d18O Ge Al

1. Introduction An organic-rich finely laminated chert (locally named Al Shomou Silicilyte) acts as a hydrocarbon reservoir in the subsurface of the South Oman Salt Basin, Sultanate of Oman. Up to 400 m thick and several kilometres wide slabs of this laminated chert are entrapped in salt domes at depths of 4e5 km. The silicilyte is

* Corresponding author. Tel.: þ41 31 631 8758; fax: þ41 31 631 4843. E-mail address: [email protected] (K. Ramseyer). 1 Present address: Department of Geosciences, University of Tübingen, Wilhelmstraße 56, D-72076 Tübingen, Germany. 0264-8172/$ e see front matter Ó 2012 Elsevier Ltd. All rights reserved. http://dx.doi.org/10.1016/j.marpetgeo.2012.08.006

a prolific source rock mature for light oil and it produces light and sour oil (using massive fracturing technology) from a highporosity, low-permeability microcrystalline silica matrix, deposited around the NeoproterozoiceCambrian boundary. The economic significance and the unconventional reservoir characteristics have triggered numerous studies to elucidate the origin and distribution of this rather unusual reservoir since its discovery in the early 1990s (e.g. Al Siyabi, 2005). Comparing the silicilyte with other siliceous deposits (Amthor et al., 2005) highlights its unique nature, because all of the possible known analogues (e.g. Precambrian banded iron formation type chert; or Phanerozoic biogenic cherts e.g. Monterey Fm.; lacustrine Magadi-type chert; or hydrothermal sinters) differ either in their source of silica, i.e. biogenic, or mode of formation,

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i.e. solid crust, or transformation characteristics, i.e. replacing amorphous silica. Despite recent progress concerning the depositional framework and origin of the silicilyte (Al Rajaibi, 2011; Amthor et al., 2005; Grosjean et al., 2009; Kowalewski et al., 2009; Schröder and Grotzinger, 2007; Wille et al., 2008) the source of silica and its mode of precipitation have generally remained unresolved. Grosjean et al. (2009) point out that a specific biomarker discovered in the silicilyte was described from demospongiae (McCaffrey et al., 1994), in which case the quartz crystals might have formed diagenetically by precipitation from dissolved biogenic silica (Schieber et al., 2000). Other potential sources of silica are from volcanic activity along mid-oceanic ridges, weathering and riverine input, inorganic precipitation from seawater as proposed for Precambrian cherts (Maliva et al., 2005) or sorption on organic molecules (Mukhopadhyay et al., 2004). In this paper we apply state-of-the-art Si and O stable isotope investigations together with trace element analyses and fluid inclusion microthermometry to further constrain the source of silica and the mode of microcrystalline quartz formation.

(e.g. Terken et al., 2001; Nederlof et al., 1997) and indicate an anoxic, sulphur-rich, stagnant and highly saline water body during the deposition of the silicilyte. In addition, the presence of X-branched compounds and Dinorhopane in the saturated hydrocarbons, in both source rocks and oils, is attributed to an origin from extinct heterotrophic bacteria and chemautotrophic bacteria, respectively (Kowalewski et al., 2009; Terken et al., 2001; Höld et al., 1999; Thiel et al., 1999; Nederlof et al., 1997).

2. Tectonic setting, stratigraphy and depositional environment

4. Results

3. Methods The study is based on core material provided by Petroleum Development Oman from 10 wells in two oil fields, a well in the central deep basin, and one on the eastern flank (Fig. 1). A total of 8 shale samples are from the Thuleilat Member and 25 silicilyte samples from the Al Shomou Member. A series of optical, isotope, geochemical and physical methods were applied to unravel the source(s) of silica and the physicochemical conditions of quartz precipitation. The detailed descriptions of the methods are given in Appendix A.

4.1. Matrix description The present-day outlines of the NE-elongated salt basin in south Oman (Fig. 1) are interpreted as regional structural trends created by NW/SE-oriented transtension during Ara depositional times (e.g. Allen, 2007). Uplifted blocks became sites of carbonate deposition, whereas basinal transtensional depressions were overlain with black shale and silicilyte (Amthor et al., 2005). Evaporites blanketed both basinal areas and uplifted blocks. The Athel Formation is one of five formations comprising the subsurface Ara Group (Amthor et al., 2003, 2005; Forbes et al., 2010) (Fig. 2). The Athel Formation (Fig. 2) includes at the base the 375 m thick Al Shomou Silicilyte Member and at the top the 140 m thick Thuleilat Shale Member. This basinal sequence overlies the 115 m thick organic rich shales of the U Formation, which is correlated with the U Carbonate Member of the A4 cycle on the adjacent carbonate platform (Amthor et al., 2005; Schröder and Grotzinger, 2007; Forbes et al., 2010). A globally correlative negative carbon isotope excursion at the base of the U Formation together with the numerical age of an ash bed at the base of the UCarbonate dated at 541  0.13 Ma (Bowring et al., 2007) reveal that the sequence was laid down just above the EdiacaraneCambrian boundary (Amthor et al., 2003, 2005; Schröder and Grotzinger, 2007). The available high-precision UePb age dates and their extrapolation bracket the deposition of the Ara Group between ca. 547 and 538 Ma (Bowring et al., 2007) and allow to estimate the average duration of carbonate/evaporite sequences to about 1.2e1.3 my. Using this average duration, the ages of the individual Ara sequences can be estimated quite precisely. Given the remaining uncertainties in terms of the stratigraphic correlation of the U and Athel Formations, a duration of 2 my (542e540 Ma) is considered to be the best estimate for the A4 sequence covering both the U and Athel Formations (Forbes et al., 2010). The Al Shomou Silicilyte was deposited during a transgressive to highstand systems tract in the deepest depressions of a segmented basin with topographic lows and highs flanked by carbonate platforms (Fig. 1). Limited detrital input reached the basin from the flanks and the silicilyte formed mainly in a stratified water mass. Recent studies by Schröder and Grotzinger (2007) and Wille et al. (2008) on redox sensitive trace elements, and by Grosjean et al. (2009) and Kowalewski et al. (2009) on source rock characteristics corroborate earlier studies on organic biomarkers in the silicilyte

Microscopy, SEM, TEM and XRD investigations (Amthor et al., 2005; Al Rajaibi, 2011) reveal the finely laminated and porous aspect of the silicilyte with a mean lamina thickness of ca. 20 mm, consisting of predominantly organic matter-rich and microcrystalline quartz-rich layers, respectively (Fig. 3A). The laminae are generally continuous on a thin-section scale but pinching-out was also observed. Moreover, they show compactional features and syndepositional deformation (Fig. 3B) indicative of soft sediment conditions during deposition and early stage diagenesis. Petrographic evidence suggests that compaction has reduced the initial sediment thickness by at least 50% of which ca. 30% is mechanical compaction (Fig. 3C). The silicilyte comprises a small quantity (<5%) of detrital components such as silt- and fine-sand sized quartz, K-feldspar, rock fragments and white mica or illite. The origin of the latter is unknown but a combined detrital/authigenic source is likely. Thinsection analyses by Amthor et al. (2005) document abundant amorphous organic matter dispersed in the silicilyte. Authigenic idiomorphic quartz crystals are the main component of the silicilyte (Fig. 3D,E). Other diagenetic phases are pyrite, apatite, dolomite, magnesite and barite. Moreover, the low crystallinity index of quartz (CI ¼ 1e4; Amthor et al., 2005) in the silicilyte is comparable to chert (CI ¼ <1e4, Murata and Norman, 1976) indicating a similar small size of the coherent diffraction domain, which is often related to the crystallite size (Amthor et al., 2005; Murata and Norman, 1976). 4.2. Geochemistry In order to determine the effect of aluminosilicates on the d30Si and d18O values in quartz and indications of the source of silica, the element concentrations were analysed by XRF, LAeICPeMS and ICPeOES. However, the micrometre crystal size of quartz and white mica and their intimate intergrowth precluded an in-situ determination of Al and Ge in individual quartz crystals. 4.2.1. Trace elements in bulk silicilyte Semi-quantitative XRF-analyses (Table 1) of eight bulk silicilyte samples show a clear covariant relationship of Al and K (r2 ¼ 0.996) indicating a dominant common source of both elements. Possible

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Figure 1. Location map showing the position of the salt basins and the structural and palaeogeographic setting of the South Oman Salt Basin during deposition of the Ara Group at the PC/C boundary. The palaeogeography during Athel Formation deposition is dominated by two carbonate platform domains surrounding a deep anoxic marine basin. Inset map shows the locations of wells used for this study. Labels A and B denote the relative positions of the stratigraphic columns shown in Figure 2.

sources are K-feldspar and micas. K-feldspar as additional source can be excluded because XRD shows no correlation with micas. Moreover, the molar K/Al-ratio of 0.20 is lower than the ratio for white mica or illite (0.25e0.35) but clearly higher than for smectite (<0.01). Therefore, a mixed-layer illiteesmectite with a SiO2/Al2O3   et al., 1986) is the most ratio of 2.5 (analysis 1M6 from Srodo n probable sheet silicate present.

Element analyses by LAeICPeMS on bulk silicilyte included measurement of 7Li, 25Mg, 27Al, 29Si, 31P, 32S, 39K, 57Fe, 73Ge, 75As, 85 Rb, 95Mo, 121Sb, 139La, 140Ce and 146Nd. All elements returned significant concentrations (see Appendix B for data of individual spot analyses and element-specific average LODs for a 60 mm crater). Cross-plots (Fig. 4) clearly document three groups of elements which correlate positively (r2 ¼ 0.98e0.75, n ¼ 47) with either Al

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Figure 2. Ara Group stratigraphic columns and nomenclature. The red line indicates the position of the PC/C boundary (cf. Bowring et al., 2007). For locations of columns A and B see Figure 1.

(K, Mg, Rb, Fig. 4A), P (La, Ce, Nd, Fig. 4B), Fe (S, As, Sb) or Ge (Mo, Fig. 4C) indicating a specific source of origin such as mica, apatitee phosphorite, pyrite or other complex sulphides, respectively. In contrast, in the case of Si, where the major contributor is quartz, the correlations with classical trace elements i.e. Ge (Fig. 4C) and Li (Appendix B), are negative and thus not incorporated in quartz. Therefore, the amount of Ge in quartz is likely <0.3 mg g1 (Appendix B, Fig. 4C), or expressed as molar Ge/Si ratio <0.25  106. This view is supported by the positive correlation of Ge with Mo and the occurrence of a few high-Ge concentration spot analyses in the silicilyte (see Appendix B). Semi-quantitative ICPeOES analyses (Table 1) document a drastic decrease of Al and K in the purified product. Using the AleK correlation from ICPeOES results, the Al concentration at zero K is 125 mg g1.

biological Si polymerization and preferential light 28Si(OH)4 adsorption on clay surfaces (Delstanche et al., 2009; Mortlock and Froelich, 1987; Scribner et al., 2006). Silicilyte d30Si bulk rock values therefore represent a mixture of isotopically lighter clay and isotopically heavier end member Si from which the microcrystalline quartz emerged. Assuming that all Al originates from the detrital clay fraction and using a SiO2/Al2O3 ratio of 2.5 (mixed-layer illitee   et al., 1986) all samples are smectite, analysis 1M6 from Srodo n below 10% Si contribution from clay except sample W-1, 1154 m which has a maximum of 11% Si from its clay fraction. No correlation between Al concentrations and d30Si values can be seen (r2 ¼ 0.008). Using the mean d30Si of 1.5& for modern clays (Fig. 5A), the calculated d30Si compositions of the microcrystalline quartz are, within long-term reproducibility of 0.23& which reveals that the clays have no influence on the d30Si of microcrystalline quartz.

4.2.2. Silicon isotopes in bulk silicilyte The d30Si signatures from eight silicilyte samples are presented in Figure 5 and Table 2. The range from 0.36 to 1.24 is narrow compared to other modern reservoirs. Being isotopically heavy in their d30Si values, these samples are close to modern aqueous Si reservoirs, which have heavy Si isotope signatures due to fractionation during

4.2.3. Oxygen isotopes in quartz Stable oxygen isotope ratios were determined in bulk and chemically purified, i.e. dissolution of non-quartz phases (Appendix A), silicilyte, and fracture filling quartz (Table 3). The range of d18O in the bulk samples varies between 24.3 and 28.1& (V-SMOW) and in the chemically purified between 26.6 and 27.7&. The observed most

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Figure 3. (A, B): Core photographs from well W2 illustrating the finely laminated nature of the silicilyte, in places displaying compacted early fractures (B). Scale bar is 4 cm (C): Planelight thin-section photomicrograph showing a detrital quartz grain in the well-laminated silicilyte (well W2). Note the laminae below the grain are deformed due to the impact (“dropstone”) and following compaction, whilst the laminae covering the grain show onlap and wedge-shaped structures typical for embedding on an uneven surface. Scale bar is 200 mm. (D): SEM image illustrating the texture of a highly porous silicilyte and the sub-microcrystalline nature of quartz, well W1. Scale bar is 10 mm. (E): Close-up of (D) illustrating the euhedral habit of the quartz crystals and the nature of the intercrystalline porosity. Scale bar is 1 mm. Figure 3 reproduced from Amthor et al. (2005) with permission of GeoArabia.

positive d18O value, i.e. 28.1& (V-SMOW) is very close to the highest values reported for Cambrian marine cherts (i.e. 29&; Knauth and Lowe, 1978; Robert and Chaussidon, 2006). The d18O value for later crystallized idiomorphic quartz varies between 19.4 and 21.2& (V-SMOW). 4.3. Fluid inclusion microthermometry Two consecutively-formed primary fluid inclusion populations I and II, i.e. trapped during crystal growth, were observed along the crystal growth zones of fracture-filling quartz. Population I contains only single-phase fluid inclusions formed prior to population II and shows high salinity containing NaCI and other chlorides. The second population comprises coexisting: i) aqueous, nonfluorescent two-phase inclusions IIa (i.e. liquid and gas phase at room temperature) with a minimal trapping temperature of 68e 83  C containing a highly saline fluid of NaCI and other chlorides; and ii) a two phase population Ilb (i.e. aqueous and hydrocarbon phase) which are fluorescent, and hence contain hydrocarbons in

addition to the water phase. Thus, population IIb indicates the presence of early hydrocarbons. 5. Discussion The origin of the oil-bearing microporous silicilyte, i.e. its mode of formation, has an important economic impact as the distribution of the organic substance, now mostly hydrocarbons, is intimately related to the formation of the silicilyte (Amthor et al., 1998, 2005). In the following discussion we will focus on the source of silica in the silicilyte and its mode of formation in an anoxic, sulphur-rich, stagnant and highly saline depositional environment (Nederlof et al., 1997; Terken et al., 2001; Amthor et al., 2005; Schröder and Grotzinger, 2007; Wille et al., 2008; Kowalewski et al., 2009). 5.1. Origin of silica The modern isotopically heavy and variable Si isotope composition of seawater is a result of preferential light Si isotope uptake

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Table 1 Analytical results of bulk rock Al, Si and K determinations by XRD and Al and K determinations by ICPeOES of purified quartz (1 mg g1 ¼ 104 wt. %). Well

W-1 W-2 W-2 W-3 W-4 W-4 W-4 W-7

Depth (m)

1154.0 4203.1 4283.2 4757.0 4426.7 4430.5 4433.2 2690.2

XRF

ICPeOES

Al2O3 (wt.%)

SiO2 (wt.%)

K2O (wt.%)

Al (mg g1)

K (mg g1)

4.33 1.33 1.87 3.13 0.77 1.15 0.90 0.76

79 93 94 92 97 96 93 97

0.789 0.239 0.326 0.545 0.128 0.178 0.144 0.112

n.d. 670 760 2000 1400 1500 1500 1500

n.d. 47 80 197 129 136 126 220

by diatoms and the continuous supply of isotopically heavy river water (De La Rocha et al., 2000; Cardinal et al., 2005; Reynolds et al., 2006; Georg et al., 2009a, 2009b; Engstrom et al., 2010; Ding et al., 2004). Today, Si has a residence time of approximately 9e15 kyr (Tréguer et al., 1995; Georg et al., 2009a). In the absence of a diatom control during the Neoproterozoic/Cambrian transition,

the concentration and distribution of Si within the oceans would have been more conservative, with homogeneous Si concentrations around 1000 mM possibly up to saturation levels of amorphous silica (Siever, 1991). Compared to the modern ocean, where the isotopically heavy Si inventory is controlled by diatom productivity, dissolution and sedimentation, the Precambrian Si cycle would have been controlled by inorganic processes like Si co-precipitation and adsorption on mineral surfaces or organic matter (Siever, 1991) resulting in a relative enrichment of 30Si(OH)4 in the remaining water. Therefore it is reasonable to assume that Neoproterozoic seawater had a homogeneous, heavy Si isotopic signature, even though a strong biosiliceous output was missing (Van den Boorn et al., 2007, 2010). The wide spatial distribution and large range in depth of the silicilyte samples together with their narrow Si isotopic composition of 0.36e1.24& imply that the Si source of the silicilyte was invariant, pointing to an isotopically homogenous reservoir enriched in isotopically heavy Si, which was most likely seawater. In addition to seawater, a number of other potential Si sources can be considered for the Neoproterozoic and Early Cambrian chert formation.

Figure 4. Cross-correlations of geochemical parameters of silicilyte samples. A) Cross plot of Al vs. Ti, Mg, and K, B) Cross plot of P vs. La, Ce, and Nd, and C) Cross plot of Ge vs. Mo and SiO2.

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Figure 5. (A): d30Si composition of main modern Si reservoirs. Data from Douthitt (1982), Ding et al. (2004), De La Rocha et al. (2000), De La Rocha (2003), Basile-Doelsch et al. (2005), Cardinal et al. (2005), Ziegler et al. (2005a, 2005b), Reynolds et al. (2006), Fitoussi et al. (2009), Georg et al. (2009a, 2009b), Opfergelt et al. (2010), Engstrom et al. (2010), Wille et al. (2010) and Hendry et al. (2010). (B): d30Si composition of Early Cambrian silicilyte samples in comparison with Archean chert samples. Figure adopted and modified from Van den Boorn et al. (2007, 2010). Table 2 d29Si and d30Si silicon isotope data from bulk silicilyte samples and corrected for the presence of clay minerals (d30Si-corr). Well

Depth (m)

d29Si (&)

d30Si (&)

d30Si-corr (&)

W-1 W-2 W-2 W-3 W-4 W-4 W-4 W-7

1154.0 4203.1 4283.2 4757.0 4426.7 4430.5 4433.2 2690.2

0.26 0.44 0.51 0.32 0.23 0.31 0.18 0.41

0.67 0.94 1.1 0.65 0.55 0.65 0.31 0.81

1.02 1.03 1.24 0.85 0.59 0.72 0.36 0.86

5.1.1. Continental origin Chemical continental weathering and its subsequent riverine discharge contribute 80% of the total dissolved Si entering modern oceans (Tréguer et al., 1995). Silicon signatures of rivers vary greatly both in concentration and isotopic composition, depending on the rate of physical and chemical erosion, Si adsorption and biological Si cycling during riverine transport (Georg et al., 2006, 2007, 2009a, 2009b; De La Rocha et al., 2000; Ziegler et al., 2005a). On average, the Si isotopic range of both modern oceanic and meteoric waters is very similar to d30Si values of the silicilyte, which does not allow the estimation of a direct influence of

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Table 3 Stable oxygen isotopes in bulk and chemically purified silicilyte and fracture filling quartz. Al2O3 data are from Table 1. Well

Depth (m)

d18OV-SMOW (&)

Al2O3 (wt%)

Occurrence

W-1 W-1 W-2 W-2 W-2 W-2 W-2 W-2 W-2 W-3 W-3 W-3 W-3 W-4 W-4 W-4 W-4 W-4 W-4 W-4 W-4 W-4 W-4 W-4 W-4 W-4 W-4 W-4 W-4 W-7 W-7 W-7 W-7

1154 1154 3983.3 4203.1 4203.1 4203.1 4283.2 4283.2 4283.2 4757.0 4757.0 4757.0 4757.0 4426.7 4426.7 4426.7 4426.7 4426.7 4426.7 4430.5 4430.5 4430.5 4430.5 4430.5 4430.5 4433.2 4433.2 4433.2 4433.2 2690.2 2690.2 2690.2 2690.2

27.27 27.52 19.40 25.50 26.90 26.60 24.30 27.00 26.80 26.43 26.70 27.10 26.80 25.70 26.10 25.90 26.10 26.90 27.20 25.80 26.50 27.40 27.20 20.80 21.20 26.10 26.60 26.70 26.60 28.10 28.07 27.60 27.70

4.3 4.3 n.d. 1.3 0.067 0.067 1.9 0.076 0.076 3.1 3.1 0.2 0.2 n.d. n.d. n.d. n.d. 0.14 0.14 1.2 1.2 0.15 0.15 n.d. n.d. 0.9 0.9 0.15 0.15 0.76 0.76 0.15 0.15

Silicilyte, bulk Silicilyte, bulk Fracture filling quartz Silicilyte, bulk Silicilyte, bulk, purified Silicilyte, bulk, purified Silicilyte, bulk Silicilyte, bulk, purified Silicilyte, bulk, purified Silicilyte, bulk Silicilyte, bulk Silicilyte, bulk, purified Silicilyte, bulk, purified Silicilyte, coarse crystals Silicilyte, coarse crystals Silicilyte, fine crystals Silicilyte, fine crystals Silicilyte, bulk, purified Silicilyte, bulk, purified Silicilyte, bulk Silicilyte, bulk Silicilyte, bulk, purified Silicilyte, bulk, purified Fracture filling quartz Fracture filling quartz Silicilyte, coarse crystals Silicilyte, coarse crystals Silicilyte, bulk, purified Silicilyte, bulk, purified Silicilyte, bulk Silicilyte, bulk Silicilyte, bulk, purified Silicilyte, bulk, purified

meteoric water on the sedimentary basin via Si isotopes alone (Fig. 5A). An average silicic acid concentration of 150 mM was estimated for modern rivers (Tréguer et al., 1995), with reported concentrations of up to 280 mM for rivers within the Bengal Basin (Georg et al., 2009a). Mass balance calculation reveal that w1/3 of the recent GangeseBrahmaputra Si flux is needed to supply the Si for silicilyte formation (Georg et al., 2009b) (see Appendix). Such a large fluvial input is not observed during the climate aridity at the time of Ara sedimentation (Mattes and Conway Morris, 1990; Schröder et al., 2003; Al Rajaibi, 2011). 5.1.2. Hydrothermal origin A significant hydrothermal Si input as suggested for the Early Cambrian black chert succession on the Yangtze Platform (Chen et al., 2009) can be excluded for the silicilyte, since the d30Si values are high compared to cherts sourced from hydrothermal venting (Fig. 5B). These cherts have identical or lower Si isotopic signatures than their parent quartz-oversaturated hydrothermal fluid because precipitation favours the light isotopes to be incorporated in the precipitate (Van den Boorn et al., 2007, 2010). Moreover, both the relatively flat REE/PAAS pattern with no positive Eu anomaly (Amthor et al., 2005) and the molar Ge/Si ratio of <0.25  106 and 0.8  0.65  106 in finely crystalline quartz and bulk silicilyte, respectively, show no indication of any hydrothermal input (Kamber and Webb, 2001; Froelich et al., 1992; Hamade et al., 2003). It should be noted, however, that the Ge/Si ratio of both the microcrystalline quartz as well as the bulk silicilyte strongly depends on early diagenetic redox conditions whereby a reducing environment favors the decoupling of Ge from Si (King et al., 2000; Hammond et al., 2000; Scribner et al., 2006) and incorporation of Ge in other phases, as shown here.

5.1.3. Biogenic origin The possible influence of sponges on the Si cycle was evaluated by chemical and isotopic mass balance calculations (Appendix C). While the chemical mass balance modelling does not rule out sponges as the Si source, the heavy isotopic values preclude sponge opal as Si source (see Appendix C for details). This result is further supported by (i) the anoxic environment on the silicilyte basin floor which does not allow metazoan life such as sponges and (ii) the lack of any evidence of siliceous sponges on the adjacent platforms which rules out input of sponge spicules into the basin. In conclusion, the measured range of d30Si in the silicilyte is best explained by a seawater source. A large direct influence of meteoric waters on the silicilyte formation cannot be excluded but is unlikely. Biogenic sources, i.e. siliceous sponges, and/or hydrothermal fluids as major determinant sources can be ruled out. 5.2. Mode of silicilyte formation The relationships between Precambrian oceanic silica concentrations and removal processes are not yet fully understood (e.g. Maliva et al., 2005). The concentration of silica in Neoproterozoic seawater was likely higher than at present due to the lack of silicasecreting planktonic organisms. A higher concentration of dissolved silica in the ocean prior to Middle Cambrian times was proposed by Siever (1991) based on saturation values of different silica phases, and is supported by the observation that early diagenetic cherts formed abundantly in peritidal marine environments (Maliva et al., 1989, 2005). In addition, the overall higher chemical weathering rates of continental silicate rocks in Cambrian times compared to the present is a further indication of increased input of silica from the continents and thus likely a higher concentration in the seawater (Racki and Cordey, 2000; François et al., 1993). This overall higher rate of silicate weathering from Late Neoproterozoic till the end of Cambrian is further supported by the presence of Transgondwanan Supermountains resulting from the assembly of Gondwana (for more details see Campbell and Allen, 2008; Campbell and Squire, 2010). Formation of silica-gel or opal-A mediated by microbes is well known (cf. Toporski et al., 2002; Yee et al., 2003; Konhauser et al., 2003; Mukhopadhyay et al., 2004; Posth et al., 2008; Jones et al., 2008; Ferris and Magalhaes, 2008; Budakoglu, 2009). However, the microbes either act as templates for nucleation and precipitation or are rapidly mineralized by opal A and not by a flexible gel (Jones et al., 2008). Moreover, laboratory experiments by Amores and Warren (2007) clearly document that silicification of microbial mats requires oversaturation with respect to amorphous silica. Hence, the process generating colloidal silica or a silica gel is independent of the microbial mats. Thus, if the silica concentration in the surface water was to be high enough to mineralize microbes, then the microbial mats on the Birba platform (Schröder, 2000) should also be silicified, and not only those in the deep basin where silicilyte formed. Furthermore, no indication exists so far that syndepositional or early-diagenetic cherts formed abundantly in peritidal environments of Ara Group carbonates (e.g. Schröder, 2000; Schröder et al., 2003). This indicates that the surface seawater was likely undersaturated with respect to amorphous silica. On the basis of textural and geochemical data, and the reconstruction of the depositional basin for the Al Shomou Silicilyte (Amthor et al., 2005), the silicilyte was formed in water quite different from that where contemporaneous carbonate formed. Hence, Amthor et al. (2005) proposed a model with a stratified water column, in which the surface waters were the site of organic productivity and platform carbonate deposition, whereas the deep water at or below the thermocline/chemocline were the site of silica-gel formation and bacterial mat growth.

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Table 4 Temperature estimates (in  C) of silicilyte matrix and fracture filling quartz based on variable seawater d18OV-SMOW values. Stable oxygen isotopic values are low, median and high values of data from Table 3 (after Friedman and O’Neil (1977) temperature dependence of the isotopic composition in quartz). Origin

Quartz d18OV-SMOW (&)

Silicilyte matrix, low d18O Silicilyte matrix, median d18O Silicilyte matrix, high d18O Fracture-filling quartz, low d18O Fracture-filling quartz, median d18O Fracture-filling quartz, high d18O

25.5 26.8 28.1 19.4 20.8 21.2

Estimated temperature ( C) for a seawater d18OV-SMOW of 0

1.4

3

4.5

5

6

74 67 59 118 106 103

65 59 52 106 95 93

57 50 44 94 84 81

49 43 37 83 74 72

46 41 35 80 71 69

42 36 31 73 65 63

The presence of organic biomarkers for chemoautotrophic bacteria and for stagnant highly saline conditions as well as redoxsensitive trace elements such as Mo in the silicilyte clearly indicate anoxic conditions and a stagnant highly saline water body below a certain depth in the basin (Amthor et al., 2005; Kowalewski et al., 2009; Schröder and Grotzinger, 2007; Wille et al., 2008). This high salinity is likely the result of sub-aqueous evaporite dissolution such as recently observed in the Bannock-, Tyro-, Ostro- and Scirocco-Basins in the eastern Mediterranean (cf. De Lange et al., 1990; Cita, 2006) and the Orca Basin in the northeastern Gulf of Mexico (cf. Pilcher and Blumstein, 2007). Thus, the drastic increase in salinity in the silicilyte basin at a certain depth is most likely due to dissolution of Ara salt atop and adjacent to the edges of transtensional depressions on the Athel deep seafloor. Besides forming a chemocline towards the shallow-marine seawater, this high salinity dramatically depresses the solubility of amorphous silica to w40% of the value in seawater (Chen and Marshall, 1982). In essence, the drastic increase in salinity is the mechanism to supersaturate amorphous silica, a prerequisite to form a silica-gel by polymerization and gelification (Iler, 1979). Moreover, this high salinity and a pH 7, as observed in the majority of the recent brine pools (e.g. De Lange et al., 1990; Cita, 2006), both support the polymerization of Si(OH)4 to form spherical particles which aggregate into three-dimensional gel networks (Iler, 1979) in the highly saline brine below the chemocline. These silica gel flakes will settle down because of their higher density, and thus cover chemautotrophic bacterial mats on the basin floor, forming a quasi lamination with layers of high bacterial- or silica-concentration. Such a genesis is further supported by the typical textural characteristics of the silicilyte, e.g. the slightly undulose lamination of predominantly organic rich and finely crystalline quartz layers, the discontinuous nature of most of the laminae with a mean thickness of ca. 20 mm, or the random presence of laminae with detrital grains (Fig. 3C). All of these textural characteristics indicate a noncontinuous process of deposition. The model of Amthor et al. (2005), therefore, should be modified insofar as the cause of silica gel formation is not the presence of bacterial mats living at the chemocline but rather the effect of silica oversaturation and polymerization at high salinity. Due to the low silica concentration in present-day seawater as a result of extraction by diatoms, no recent analogue exists in the brine pools known from the eastern Mediterranean or the northeastern Gulf of Mexico (cf. De Lange et al., 1990; Cita, 2006; Pilcher and Blumstein, 2007).

The oxygen isotopes of quartz are used to cross-check these temperature estimates. Using the median d18O value of the silicilyte matrix and a temperature of <45  C then the seawater d18O must be less than 4.5& (Table 4). The fact that this value is within the range of Cambrian seawater, as modelled by Jaffrés et al. (2007), further supports the proposed low-temperature lithification of the silicilyte. The lithification process may either be direct dissolutione reprecipitation of silica gel to microcrystalline quartz or a two-step reaction through an intermediate phase such as opal-CT. No petrographic indication of an intermediate phase such as opal-CT lepispheres was found, however. Nevertheless, the high Al concentration in the microcrystalline quartz, i.e. <125 mg g1, is supporting evidence of a non-crystalline precursor phase as documented by Götze et al. (2009) for chalcedonyemacrocrystalline quartz pairs in agates.

5.3. Significance of lamination

6. Conclusions

As stated earlier, the discontinuous nature of the lamination reflects a non-continuous process of bacterial mat growth on the basin floor and deposition of silica-gel flakes. A rough estimate of the mean number of laminae formed per year can be obtained by taking the 375 m of silicilyte strata (Amthor et al., 2005), a constant sedimentation rate during A4C cycle, i.e. for the Thuleilat Mbr., Al

Combined evidence indicates that the Al Shomou Silicilyte formed by primary silica precipitation (as opposed to replacement of a carbonate or other mineral precursor) in a highly saline, anoxic, basinal environment. It is therefore regarded as a primary chert (Maliva et al., 2005) that formed at the Neoproterozoic/Cambrian boundary.

Shomou Mbr. and U-Shale Mbr. or the total of 630 m section, and the estimated duration of 1 Ma for the A4C cycle (Forbes et al., 2010). Based on these data about 32 laminae of 20 mm would have been formed yearly. Thus, these laminae are not annual cycles but cycles of shorter duration. The origin of these cycles is ambiguous, as possible causes are fluctuating silica concentration in the seawater triggered by climatic effects such as periodic rainfall, lunar forced effects, i.e. spring tidal re-mixing of the water body or diagenetic self-organization (Murray et al., 1992). 5.4. Timing and mode of lithification of silicilyte Several lines of independent evidence indicate a lowtemperature, i.e. early time, of lithification of the silicilyte: (i) The observed compactional loss (Fig. 3C) implies a burial depth of ca. 750 m assuming compactional loss similar to shale (Schlumberger PetroModÒ 1D freeware; Hantschel and Kauerauf, 2009). This corresponds to ca. 45  C burial temperature (pers. comm. PDO). (ii) The presence of all-liquid single phase fluid inclusions indicates a minimal temperature of 50  C (Goldstein and Reynolds, 1994). (iii) Taking the occurrence of reworked silicilyte pebbles in the ca. 5 Ma years younger Irad Mbr. sandstone of the Karim Fm. (unpublished data PDO), then the inferred upper temperature is <45  C.

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Newly acquired geochemical data support ambient seawater rather than hydrothermal or biogenic sources (e.g. sponges) as the primary source of the silica. The molar Ge/Si ratio and the Ge concentration of the authigenic microcrystalline quartz of bulk silicilyte are indicative of an early diagenetic decoupling of Ge and Si with preferential incorporation of Ge into other phases. Therefore, the molar Ge/Si ratio is not indicative of the seawater ratio but rather of early diagenetic reactions under anoxic conditions. Silica precipitation from seawater was the result of a dramatic increase of salinity through dissolution of sub-aqueously exposed evaporites, resulting in a prominent depression of amorphous silica solubility in seawater. Resulting saturation caused the formation of silica-gel flakes. Sedimentation of these flakes onto bacterial mats covering the basin floor generated an undulose and discontinuous lamination of silica-rich and bacterial mat-rich layers. Lithification of the silicilyte occurred early at temperatures below 45  C in the presence of water strongly depleted with respect to 18O. Acknowledgements The authors acknowledge Petroleum Development LLC and the Ministry of Petroleum and Gas of the Sultanate of Oman for permission to publish the paper. The Si isotope work was funded through Australian Research Grants DP0770820 and DP0771519 awarded to Bill Maher (UC), Michael Ellwood (ANU) and Steve Eggins (ANU). Chris Taylor is thanked for technical assistance at SUERC. We thank two anonymous reviewers for their constructive criticism which helped to improve the manuscript. Appendices AeC. Supplementary data Supplementary data related to this article can be found online at http://dx.doi.org/10.1016/j.marpetgeo.2012.08.006. References Al Rajaibi, I.M.A., 2011. Origin and Variability of the Late PrecambrianeCambrian Athel Silicilyte, South Oman Salt Basin. PhD thesis, The University of Manchester, Manchester. Al Siyabi, H.A., 2005. Exploration history of the Ara intrasalt carbonate stringers in the South Oman Salt Basin. GeoArabia 10, 39e72. Allen, P.A., 2007. The Huqf Supergroup of Oman: basin development and context for Neoproterozoic glaciation. Earth-Science Reviews 84, 139e185. Amores, D.R., Warren, L.A., 2007. Identifying when microbes biosilicify: the interconnected requirements of acidic pH, colloidal SiO2 and exposed microbial surface. Chemical Geology 240, 298e312. Amthor, J.E., Smits, W., Nederlof, P., Frewin, N.L., Lake, S., 1998. Prolific Oil Production from a Source Rock e the Athel Silicilyte Source-Rock Play in South Oman. American Association of Petroleum Geologists Annual Convention, 122 pp. Amthor, J.E., Grotzinger, S., Schröder, S.A., Bowring, J., Ramezani, M., Martin, M., Matter, A., 2003. Extinction of Cloudina and Namacalathus at the Precambriane Cambrian boundary in Oman. Geology 31, 431e434. Amthor, J.E., Ramseyer, K., Faulkner, T., Lucas, P., 2005. Stratigraphy and sedimentology of a chert reservoir at the PrecambrianeCambrian boundary: the Al Shomou Silicilyte, South Oman Salt Basin. GeoArabia 10, 89e122. Basile-Doelsch, I., Meunier, J.D., Parron, C., 2005. Another continental pool in the terrestrial silicon cycle. Nature 433, 399e402. Bowring, S.A., Grotzinger, J.P., Condon, D., Ramezani, J., Newall, M., Allen, P.A., 2007. Geochronologic constraints on the chronostratigraphic framework of the Neoproterozoic Huqf Supergroup, Sultanate of Oman. American Journal of Science 307, 1097e1145. Budakoglu, M., 2009. Comparison of recent siliceous and carbonate mat developuk Lake, NE ment on the shore of hyper-alkaline Lake Van and Mt. Nemrut Sog Anatolia, Turkey. Geomicrobiology Journal 226, 146e160. Campbell, I.H., Allen, C.M., 2008. Formation of supercontinents linked to increases in atmospheric oxygen. Nature Geoscience 1, 554e558. Campbell, I.H., Squire, R.J., 2010. The mountains that triggered the Late Neoproterozoic increase in oxygen: the Second Great Oxidation Event. Geochimica et Cosmochimica Acta 74, 4187e4206. Cardinal, D., Alleman, L.Y., Dehairs, F., Savoye, N., Trull, T.W., Andre, L., 2005. Relevance of silicon isotopes to Si-nutrient utilization and Si-source assessment in Antarctic. Global Biogeochemical Cycles 19. GB20071-13. Chen, C.-T.A., Marshall, W.L., 1982. Amorphous silica solubilities IV. Behavior in pure water and aqueous sodium chloride, sodium sulfate, magnesium

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