The role of bacterial sulfate reduction during dolomite precipitation: Implications from Upper Jurassic platform carbonates

Chemical Geology 412 (2015) 1–14

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The role of bacterial sulfate reduction during dolomite precipitation: Implications from Upper Jurassic platform carbonates Andre Baldermann a,⁎, Artur P. Deditius b, Martin Dietzel a, Vanessa Fichtner c, Cornelius Fischer d, Dorothee Hippler a, Albrecht Leis e, Claudia Baldermann f, Vasileios Mavromatis a, Christian P. Stickler a, Harald Strauss c a

Institute of Applied Geosciences, Graz University of Technology, Graz, Austria School of Engineering and Information Technology, Murdoch University, Murdoch, Australia c Institut für Geologie und Paläontologie, Westfälische Wilhelms-Universität, Münster, Germany d MARUM/FB Geowissenschaften, Universität Bremen, Bremen, Germany e RESOURCES - Institute of Water, Energy and Sustainability, Joanneum Research, Graz, Austria f Institute of Technology and Testing of Building Materials, Graz University of Technology, Graz, Austria b

a r t i c l e

i n f o

Article history: Received 17 April 2015 Received in revised form 13 July 2015 Accepted 14 July 2015 Available online 17 July 2015 Keywords: Dolomite Diagenesis Platform carbonates Stable isotopes Trace elements Bacterial sulfate reduction Carbonate-associated sulfate

a b s t r a c t The early diagenetic formation of dolomite in modern aquatic environments is limited mostly to evaporitic and marine-anoxic, organic-rich sediments dominated by bacterial sulfate reduction (BSR). In such environments, bacterial activity lowers the energy barriers for the nucleation and growth of dolomite and thus promotes the formation of non-stoichiometric, highly disordered and metastable (proto)dolomite. Although the boundary conditions for the formation of modern (proto)dolomites are considered to be generally understood, the role of BSR during limestone dolomitization in ancient marine environments remains questionable. Herein, we present a study about the physicochemical conditions and processes, which led to the formation of partly dolomitized limestone and dolostone in the presence of BSR on a stable carbonate platform during the Upper Jurassic at Oker (Northern German Basin). The dolomite textures, the spatial trace elemental patterns of the dolomite and of the surrounding limestone and the results of δ18O and δ13C isotope analyses reveal that the Oker dolomite has been formed by the early diagenetic replacement of magnesian calcite precursors at temperatures between 26 °C and 37 °C. We interpret the mineralizing fluids responsible for dolomitization as pristine-marine to slightly evaporitic and reducing seawater being modified during shallow seepage reflux and/or evaporitic tidal pumping. The elevated δ34SCAS values (+17.9 to +19.7‰, V-CDT) of the Oker dolomite, compared to ambient Upper Jurassic seawater, indicate that BSR facilitated dolomite formation. For the first time, we show that a linear anti-correlation exists between decreasing carbonate-associated sulfate (CAS) contents in dolomite and increasing ordering ratio of the dolomite lattice structure, with the degree of cation order in dolomite to be given by: degree of cation order(Dol): = −0.018·CAS(Dol) + 68.3 (R2 = 0.98). This correlation implies that the CAS content of sedimentary dolomite can be used as a measure for dolomite maturity. The relationships between the ambient (paleo)environmental controls, the resultant dolomitization pathways and subsequently the structure and the composition of the precipitating dolomite are presented and discussed in relation to the stability of modern and ancient (proto)dolomites throughout burial diagenesis. © 2015 Elsevier B.V. All rights reserved.

1. Introduction Dolomite is a common rock-forming mineral in ancient carbonate platforms, but its formation in modern marine environments is rare, although the ambient seawater is supersaturated with respect to dolomite by one to two orders of magnitude (Lippmann, 1973; Machel & Mountjoy, 1986; Arvidson & MacKenzie, 1999; Warren, 2000). Numerous ⁎ Corresponding author at: Institute of Applied Geosciences, Graz University of Technology, 8010 Graz, Austria. E-mail address: [email protected] (A. Baldermann).

http://dx.doi.org/10.1016/j.chemgeo.2015.07.020 0009-2541/© 2015 Elsevier B.V. All rights reserved.

models have been developed in the last decades in order to explain the origin of sedimentary dolomite and its abundance in the geological record. Conceptual models range from direct precipitation of dolomite from Mg-rich mineralizing fluids to the early diagenetic transformation of metastable carbonate precursors to dolomite (Machel & Mountjoy, 1986; Budd, 1997; Warren, 2000). More recently, the relation between the occurrence of volumetrically significant amounts of sedimentary dolomite and variations of the ambient paleo-environmental conditions has been used successfully for the reconstruction of dolomite-forming processes in ancient marine environments (Burns et al., 2000). It has been proposed that bacterial activity and bacterial sulfate reduction

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(BSR) could play an important role during the dolomitization of ancient platform carbonates, by lowering the energy barriers that inhibit dolomite precipitation from oversaturated solutions (Burns et al., 2000; Loyd et al., 2012; Meister et al., 2013). The results of previous experimental work indicate that microbial activity helps to overcome imposed energy barriers limiting the nucleation of dolomite and thus is an essential parameter for its precipitation at low temperatures (Vasconcelos et al., 1995; Land, 1998; Warthmann et al., 2005; Deng et al., 2010). Similar effects have been suggested for ion activity, moderate to high Mg/Ca ratios, alkaline pH, high CO2− 3 high Fe2+ and dissolved sulfide concentrations, low sulfate concentrations, elevated temperature and presence of seeds (e.g., Lippmann, 1973; Land, 1980; Baker & Kastner, 1981; McKenzie, 1981; Lumsden et al., 1995; Vasconcelos et al., 1995; Arvidson & MacKenzie, 1999; Warren, 2000; Zhang et al., 2012; Geske et al., 2015). Since the development of the microbial model (Vasconcelos & McKenzie, 1997), the links between microbial activity, pore fluid geochemistry and the preferential precipitation of dolomite in micro-environments developing around the cells have been widely decoded (e.g., Warthmann et al., 2005; Bontognali et al., 2010). It has been suggested that the presence of exopolymeric substances (EPS) secreted by sulfate-reducing bacteria could play a key role in the precipitation of Ca-excess dolomite and Mg-calcite at low temperatures by enhancing the desolvation of Mg2+ aquo-complexes at crystal surface sites and thus promoting Mg2+ ion incorporation into the precipitating Ca–Mg-carbonates (Bontognali et al., 2014). Thus, conditions favorable for the recent formation of dolomite occur in evaporitic and marine-anoxic environments, in association with organic-rich sediments and sulfate-reducing bacteria, as reported for the modern sabkha environments of Abu Dhabi and the Coorong Region, South Australia (e.g., Wacey et al., 2007; Bontognali et al., 2010) as well as in the hypersaline coastal lagoon at Lagoa Vermelha, Brazil (e.g., Vasconcelos & McKenzie, 1997; Warthmann et al., 2005). In these environments, the precipitating non-stoichiometric, Caexcess dolomite is characterized by a poorly ordered structure and is thus sometimes referred to as (proto)dolomite (Warren, 2000). Subsequent “aging” processes associated with burial diagenesis, i.e. inorganic multiple recrystallization of the metastable (proto)dolomite, are currently suggested to result in a more stoichiometric and ordered type of dolomite (Burns et al., 2000; Warren, 2000). In this study, we expose the importance of BSR during the dolomitization of ancient platform carbonates. Therefore, a succession of Upper Jurassic partly dolomitized limestone and dolostone as well as cogenetic limestone was investigated, which has been deposited on a stable carbonate platform at Oker (Northern German Basin). From the results of field work, thin-section analyses, geochemical modeling, δ18O, δ13C and carbonate-associated sulfate–sulfur (δ34SCAS) isotope analyses and trace element distributions and concentrations of the platform carbonates from Oker, we reconstruct and discuss the physicochemical conditions and reaction pathways leading to the dolomitization of platform carbonates in an ancient marine-diagenetic environment. We highlight here the important role of BSR on the formation of sedimentary dolomite and show that the amount of CAS in dolomite can be used as a measure for its maturity. Herein, the links between the physicochemical conditions and the underlying dolomite-forming processes are discussed in relation to the composition and the structure of modern and ancient sedimentary dolomites and their stability during burial diagenesis. 2. Geological setting and lithostratigraphy The Langenberg section at Oker is exposed in an active quarry and is located 5 km east of Goslar in the Northern German Basin north of the Harz Mountains (Fig. 1A). This location exhibits a succession about 150 m thick of Upper Jurassic rocks (Fig. 1B). The basal part of the Langenberg section comprises oolitic grainstones and Exogyra boundstones of the Higher Coral Rag (Upper Oxfordian) that were

deposited on a stable, flat-angled carbonate platform located at subtropical latitudes (31–34°) and partly connected with the Tethys Ocean (Fischer, 1991; Thierry, 2000; Brigaud et al., 2008). The Lower to Middle Kimmeridgian rocks from Oker are characterized by a lagoonal facies of alternating argillite, mudstone and dolostone with sections containing a moderate to high terrigenous input from the adjacent Pompecki block (Fischer, 1991; Baldermann et al., 2012). Conglomerates and mudstones with Skolithos are exposed at the top of this section, which suggest a shift back to open marine conditions during the Upper Kimmeridgian (Fischer, 1991). The Cretaceous rocks that lay discordant on the Upper Jurassic succession are currently not exposed. Due to rapid subsidence of the Mesozoic sediments in the Upper Jurassic to Lower Cretaceous, the Langenberg section was buried to a mean depth of 1500–2000 m, which corresponds to a maximum burial temperature from 60 °C to 75 °C (e.g., Nollet et al., 2005; Voigt et al., 2006; Baldermann et al., 2012). Ongoing with the development of the Harz Thrust Fault in the Upper Cretaceous and a related fast uplift of the Harz Mountains, the Langenberg section was inverted and now dips at an angle of 50–70° to the south (Fig. 1B). Thus, the subsurfaces of the beds are exposed. A precise biostratigraphic correlation of the Langenberg sedimentary rocks with the standard sub-boreal ammonite zonation is problematic, because index fossils are rare at Oker (Fischer, 1991). Mudroch (2001) reported sediment deposition ages from 158 to 153 Ma (Upper Oxfordian to Upper Kimmeridgian) based on combined Rb/Sr age calculations of biogenic apatite and the biolithostratigraphic classification of Fischer (1991). 3. Material and analytical methods 3.1. Field work and samples Geological mapping of the Langenberg section was carried out during three field campaigns in June 2012, August 2013 and May 2014. Six sections with a total vertical thickness of ~250 m were logged and studied bed by bed at a minimum resolution of two centimeters. The main types of limestone microfacies were classified and interpreted in accordance with the scheme of Tucker and Wright (1990) and Flügel (2004). The dolomite rock texture classification is based on Sibley and Gregg (1987) and the nomenclature of the dolomite types is derived from Reinhold (1998). For petrographic, (micro)textural, mineralogical and (isotope)geochemical characterization of the mostly calcareous to marly sedimentary rocks, oriented specimens (66 in total) were collected from the entire Langenberg section (Fig. 2). Thin sections were prepared in order to study the depositional facies and diagenetic alteration by optical and scanning electron microscopy. 3.2. X-ray diffraction X-ray diffraction (XRD) analyses were conducted for mineral identification and quantification using a PANalytical X'Pert PRO diffractometer equipped with a Co-target tube (operated at 40 kV and 40 mA), Scientific X'Celerator detector, 0.5° antiscattering and divergence slits, spinner stage, primary and secondary soller, and automatic sample changer. The samples were first mixed with 10 wt.% ZnO as internal standard and then finely-ground in a McCrone micronizing mill for 8 min. Subsequently, randomly oriented preparations were made using the top loading technique and examined over the range of 4–85° 2θ with a step size of 0.008° 2θ/s and a count time of 40 s/step. Mineral quantification was obtained by Rietveld analysis of the XRD patterns using the PANalytical X'Pert HighScore Plus Software. Dolomite stoichiometry was calculated using the expression:

mol% CaCO3 ¼ 333:33  dð104Þ −911:99 (e.g., Royce et al., 1971; Lumsden & Chimahusky, 1980), which is based

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Fig. 1. A: Geological map of the Northern German Basin (modified from Voigt et al.(2006)) with the location of the Upper Jurassic Langenberg section at Oker in the north of the Harz Mountains. B: Cross-section of the Langenberg section including the stratigraphic head beds of Fischer (1991). Note that the section was tectonically inverted. Person for scale is 1.7 m tall. (For a colored version of this figure, the reader is referred to the web version of this article.)

on the linear relationship between dolomite composition and the position of the d(104)-peak of dolomite. Reeder and Sheppard (1984) demonstrated that this equation can lead to inaccuracies of 1 and 3 mol% CaCO3 for Ca-excess and stoichiometric dolomite, respectively. In the present study, the sample displacement was corrected to the peak positions of ZnO, resulting in a reproducibility of the dolomite d(104)-peak position within ±0.02° 2θ, equivalent to an error of 0.6 mol% CaCO3 in dolomite. The ordering ratio in dolomite was derived from the intensity ratio of the d(105)/d(110) reflections (Goldsmith & Graf, 1958; Füchtbauer & Goldschmidt, 1965). This approach is semi-quantitative and its

accuracy depends mainly on the particle orientation and the particle size, which result in an overall error of ±5%. 3.3. Electron microprobe The analyses of the geochemical composition of limestone, dolomitized limestone and dolostone samples were performed on a Jeol JXA8200 SuperProbe Electron Probe Microanalyzer (EMPA), operated at the Montanuniversität Leoben (Austria). Due to the beam sensitive nature of the investigated material the analytical conditions were as

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Fig. 2. Lithostratigraphic profile of the Langenberg section, showing the position (from left to right) of sampled intervals, the correlation with the stratigraphic head beds of Fischer (1991) and the δ18O and δ13C isotope profiles. Glt: Glauconite-bearing head beds. The location of prominent dolostone horizons and partly dolomitized limestone is highlighted in yellow color. Characteristic photographs (A to F left column) and photomicrographs (A to F right column) of polished rock chips and thin-sections (crossed nicols) of dolomitized beds are shown on the right. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

follows: 15 keV accelerating voltage, 10 nA beam current and focused beam ~1 μm in size scanning over a raster of 4 μm × 6 μm. The quantitative geochemical analyses were standardized against natural and synthetic crystals, and included the following elements with characteristic spectral lines: Ca-Kα and Mg-Kα (dolomite), Fe-Kα (almandine), MnKα (rhodonite), Sr-Lα and S-Kα (celestine) and Na-Kα (albite). The counting times for (i) Ca, Mg, Fe, Mn and Na were 20 s on peak and 10 s on the background position on each side of the peak, and for (ii) Sr and S 40 s on peak and 20 s on the background position, respectively. Only the analyses with an analytical error of b5% were included into the further consideration. The lower analytical totals are due to the presence of CO2 and porosity (see Electronic Appendix A for geochemical

composition data). The element distribution maps of 1000 × 1000 pixel resolution were collected using focused beam ~1 μm in size, 20 keV accelerating voltage, 20 nA beam current and a dwell time of 20 ms/step (see Electronic Appendix B for element distribution maps). 3.4. Stable isotopes The stable oxygen and carbon isotopic composition of carbonate minerals was measured with a ThermoFinnigan DELTAplusXP mass spectrometer equipped with a Finnigan Gasbench II at the Joanneum Research institute stable isotope laboratory (Graz, Austria). Sample (pre)-treatment is described in detail in Dietzel et al. (2009). Briefly,

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prior to the continuous-flow isotopic ratio mass spectrometric measurements, concentrated phosphoric acid was injected into the individual sample vials containing 200–400 μg of sample, and reacted with the carbonates for 2 h at 90 °C to ensure complete digestion of dolomite. The oxygen and carbon stable isotope values are reported in terms of delta notation (δ13C and δ18O) versus the international V-PDB standard. The analytical reproducibility was better than ± 0.1‰ for δ13C and δ18O, respectively. Carbonate-associated sulfate (CAS) was extracted from dolostone (n = 6) and limestone (n = 2) using the sequential leaching method described by Wotte et al. (2012). Powdered samples (91–278 g) were leached with a 10% NaCl solution in order to remove non-CAS components such as secondary sulfates. This extraction step was repeated until no further sulfate was liberated, as determined via BaSO4 precipitation. The residual rock powder was subsequently dissolved in a 25% HCl solution in order to liberate CAS from the calcite and dolomite crystal lattice. The resulting solutions were filtered through 0.45 μm cellulose nitrate membrane filters, and the dissolved sulfate was precipitated as BaSO4 by the addition of 8.5% BaCl2 solution (80 °C, 3 h). The δ34SCAS signatures of the collected BaSO4 were measured with a ThermoFinnigan Deltaplus mass spectrometer connected to a Carlo Erba Elemental Analyzer at the Westfälische Wilhelms-Universität Münster (Germany). The results are reported relative to the international V-CDT standard. The analytical reproducibility for δ34SCAS values was better than ±0.3‰. 4. Results The combination of field work and thin-section analysis provides a ~ 90 m thick lithostratigraphic profile of the Langenberg section at Oker (Fig. 2). In the following sections, the petrographic, mineralogical, trace elemental and stable isotopic results are presented for the recognized main types of limestone microfacies and dolomite, respectively (see Electronic Appendix C for stable isotopes and bulk mineralogy). The main focus is on the characterization of the Oker dolomite. 4.1. Limestone Limestone comprises ~ 85–90 vol.% of the estimated total rock volume at Oker (Fischer, 1991). Four main microfacies types were identified in the field and in thin sections (Fig. 3). 4.1.1. Oolitic grainstones and peloidal–oncolitic packstones Grayish (yellowish-brown when weathered), tabular cross-bedded, oolitic to oncolitic grainstones (Fig. 2A) dominate the first 60 m of the Langenberg section (beds 1–23). These sedimentary rocks (− 2.4 to −1.8‰ for δ18O and +2.4 to +2.6‰ for δ13C) comprise mainly of ferric ooids and oncoids that are accompanied by strictly marine bioclastic debris and small amounts of quartz (1–3 wt.%). Neomorphic calcite spar is scarce, whereas void-filling dolomite cement (16–19 wt.%) is abundant (see Section 4.2.4 for detailed description). The beds 25–53 consist of peloidal–oncolitic packstones (bed 24 is an Exogyra boundstone, see Section 4.1.2), which are intercalated with argillite and mud- to wackestones. In contrast to the segment with mostly oolitic grainstones, rocks from this section show no sedimentary structures due to intensive bioturbation of type Thalassinoides. The peloidal– oncolitic packstones (− 1.7 to − 0.3‰ for δ18O and − 2.1 to + 1.9‰ for δ13C) are composed of blackened (organic-rich and partly pyritized) peloids, ooids, oncoids and bioclastic debris (Fig. 3A and B), in addition to micrite, and high amounts of clay minerals (up to 22 wt.% of kaolinite, illite and chlorite), pyrite (0–4 wt.%) and quartz (1–14 wt.%). 4.1.2. Exogyra boundstones Exogyra boundstones (bed 24) occur as “patchy” micro-reefs in the Langenberg section, with a reef diameter ranging from 0.5 to 1.5 m (Fig. 3C and D). Their δ18O isotopic composition ranges from − 1.6 to

Fig. 3. Representative photographs (A, C, D, E, I), thin-section photomicrographs (B, F, J; crossed nicols), and BSE images (G, H) displaying the main types of limestone microfacies at Oker (oolitic grainstones are not included). A–B: Peloidal–oncolitic packstone (sample #13). C–D: Exogyra boundstone (sample #6). E–H: Mudstone–wackestone (sample #54). I–J: Floatstone–rudstone (sample #63). (For a colored version of this figure, the reader is referred to the web version of this article.)

− 1.3‰ and δ13C values vary between + 1.4 and + 1.6‰. The reef framework is strengthened by strictly marine faunal elements. The reef cavities are typically filled with ooids and pisoids as well as poorly-sorted bioclastic debris of bivalves, brachiopods, and echinoderms. Micrite is scarce. The occurrence of the Exogyra micro-reefs in

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the Langenberg section denotes the transition from an open marine, highly energetic setting to sediment deposition in a low-energy, shallow-lagoonal environment. 4.1.3. Mudstones and wackestones Grayish mudstones and wackestones are the most abundant sedimentary rocks at Oker (Figs. 2, 3E and F). The strongly bioturbated, partly glauconite-bearing, mudstones (beds 54–57 and 130–149) and wackestones (beds 71–126) consist of micrite (92–99 wt.%) and minor amounts of quartz and detrital clay minerals (b 5 wt.%). Columnar and equant calcite spar precipitated frequently in biomoldic pores and few secondary veins (Fig. 3G and H). Dolomitization is negligible. This microfacies type (− 3.4 to + 0.1‰ for δ18O, − 4.0 to + 1.4‰ for δ13C and +19.1 to +22.2‰ for δ34SCAS) exhibits a marine, low-diversity fossil association, which is dominated by Goniolina, Isastrea, and Trigonia sp. Remnants of Characeae or charophytes (beds 112–141) and Skolithos burrows (beds 147–149) rarely occur in the upper part of this unit. 4.1.4. Floatstones and rudstones Beds 131, 138, 140 and 144 (Fig. 3I and J) comprise grayish, slightly bituminous floatstones and rudstones (− 2.6 to − 0.3‰ for δ18O and − 4.8 to − 2.5‰ for δ13C). These sedimentary rocks are ~ 0.1 to 0.8 m thick, show erosive sub-surfaces and consist mainly of a few centimeter sized pebbles of reworked Kimmeridgian limestone (N96 wt.% of calcite with traces of quartz and detrital clay minerals), embedded in light gray micrite. Fischer (1991) determined the origin of the calcareous pebbles to be intraformational and suggested that their grayish to blackened color is caused by interactions with humic acids during terrestrial weathering.

4.2. Dolostone Three prominent dolostone horizons were distinguished in the Langenberg section (termed “lower dolostone”, “middle dolostone”, and “upper dolostone”, from bottom to top). Volumetrically less important and only partly dolomitized beds of oolitic–oncolitic grainstone and mudstone occur as well. The location of the dolomitized segments as well as photographs and photomicrographs of polished rock chips and thin-section of the Oker dolomite are shown in Fig. 2. The dolomitized intervals account for ~ 10–15 vol.% of the estimated total rock volume at Oker (Fischer, 1991). Four types of dolomite were identified in the field and in thin-section, which are described below. 4.2.1. Matrix dolomite-A Matrix dolomite-A is a fine-grained (10–50 μm in size), predominantly euhedral (planar-e type) to subhedral (planar-s type) dolomite, which formed at the expense of a pre-existing micrite. Type-A dolomitization encounters for b1 vol.% of the estimated total dolomite and occurs only in the beds 58–61. This type of dolomite is always found in close association with organic-rich burrows that also contain glauconite, framboidal and/or cuboidal pyrite and illite–smectite, embedded in micrite (Fig. 2C). Larger (30–50 μm sized) dolomite rhombs dominate towards the surrounding mudstone (Fig. 4A), which additionally contains partly sparitized bioclastic debris, quartz (5–7 wt.%) and detrital clay minerals (6–13 wt.%). The irregular spatial distribution of dolomite crystals in the burrow-like outlines indicates that dolomitization is incomplete (Reinhold, 1998; Rameil, 2008), explaining the strong variation in dolomite content, from 4 to 43 wt.%, at the μm- to cm-scale. The non-stoichiometric (59.7 mol% CaCO3) matrix dolomite-A contains

Fig. 4. BSE images displaying the characteristic features of the Oker dolomite. Four types of dolomite were identified using the dolomite classification scheme of Reinhold (1998). A: Matrix dolomite-A occurs in isolated patches only within intensively bioturbated, glauconitic mudstone (sample #23). Dol: dolomite. Py: pyrite. I–S: illite–smectite. B: Matrix dolomite-B from the upper dolostone (sample #50). C: Matrix dolomite-C and vein-filling calcite spar from the boundary between the upper dolostone and the underlying mudstone (sample #49). Note the association of dolomite (dark cores), low-Mg calcite (bright area surrounding the core) and high-Mg calcite (dark outermost zone), deposited in single grains (marked with arrows in the inserted image). D: Planar and limpid dolomite occurs as void-filling cement in oolitic grainstones (sample #3).

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moderate to high amounts of Na (250–700 ppm), Sr (b50-400 ppm) and S (250–650 ppm), and shows relatively low Fe and Mn contents (b 50–1600 ppm). The chemical formula of the dolomite is (Ca1.045–1.105Na0.007–0.018Sr 0–0.001) 1.056–1.120 (Mg0.841–0.923 Fe0–0.005 Mn0–0.001)0.844–0.928[(C0.998–0.999S0.001–0.002)O3]2. Matrix dolomite-A displays a δ18O isotopic composition of + 1.7 to + 1.9‰, the δ13C values range from + 0.2 to + 0.5‰ and the δ34SCAS isotopic composition is + 17.9‰. 4.2.2. Matrix dolomite-B Matrix dolomite-B forms well-developed, ~ 1.5–3.0 m thick, stratiform dolostone beds (beds 27, 62–65, and 126–129) that are easily distinguishable from the surrounding mud- and wackestones due to their noticeably lower weathering resistance and yellowish–brownish color (Fig. 2B, D and F). Type-B dolomitization accounts for ~60–70 vol.% of the dolomitized rock volume and forms medium- to fine-grained (10–80 μm in size), planar-e to planar-s shaped dolomite rhombs (Fig. 4B). In contrast to the other types of dolomite, matrix dolomite-B is an almost pure (77–97 wt.% of dolomite) dolostone, which contains minor amounts of quartz (1–7 wt.%), clay minerals (1–16 wt.%) and traces of micrite. This type of dolomite forms hypidiotopic to idiotopic rock textures with a high inter-particle porosity (40–50 vol.%) and is usually fabric-destructive, so that no microfacies features of the precursor rock are preserved. However, the fenestral fabrics, tepee structures and birdeyes in the lower and middle dolostone and remnants of algae lamination and caliche structures in the upper dolostone represent relicts of mudstone microfacies and indicate deposition in tidal flats (Pomoni-Papaioannou, 2008). The non-stoichiometric (54.0– 58.7 mol% CaCO3) matrix dolomite-B displays zoned dolomite crystals with brighter Na-, Sr- and S-rich cores and darker, partly Fe-rich, rims (Fig. 4B). Such zonation pattern in dolomite is often termed in the literature as cloudy-center-clear-rim texture (e.g., Tucker & Wright, 1990; Budd, 1997). EMP analyses revealed moderate to relatively high and variable amounts of Na (50–700 ppm), Sr (b 50–700 ppm), S (b 50–950 ppm) and low to moderate Fe and Mn contents (b50– 8550 ppm). The chemical composition of this type of dolomite is (Ca1.020–1.150Na0.001–0.018Sr 0–0.001) 1.032–1.161 (Mg0.817–0.933 Fe0–0.027 Mn0–0.006)0.819–0.937[(C0.997–1S0–0.003)O3]2. Its isotopic composition ranges from + 1.4 to + 2.9‰ for δ18O, − 0.1 to + 2.0‰ for δ13C and +19.0 to +19.7‰ for δ34SCAS. The trace elemental and isotopic composition of matrix dolomite-B is similar to that of matrix dolomite-A. 4.2.3. Matrix dolomite-C Matrix dolomite-C is a planar-s typed dolomite, which occurs in a single bed (Fig. 2E, bed 125) of ~ 10–30 cm thickness, within the Langenberg section. This horizon is located directly at the boundary between the upper dolostone and the underlying mudstone (Fig. 2). The boundary is well-defined by a sharp-erosive, lower surface to the mudstone and a more diffuse zone to the overlying dolostone. In contrast to the dolomite types A and B, the rhombic crystal shape of the dolomite grains (~ 8 wt.%) is poorly defined and recrystallized micrite, so-called (micro-)sparite, dominates (69–73 wt.%), together with quartz (7–8 wt.%) and clay minerals (12–15 wt.%). This type of dolomite reveals a high inter-particle porosity of ~ 30–40 vol.% and encounters for only ~ 1 vol.% of the total estimated dolomite. Tubular, secondary veins are frequently filled with blocky calcite spar. EMP analyses and backscattered electron (BSE) images (Fig. 4C) document that the matrix dolomite-C consists of fine-grained (4–20 μm in size) dolomite cores, (Ca0.969–1.057Na0–0.015Sr0–0.001)0.978–1.072(Mg0.904–0.985 Fe0–0.021Mn0–0.052)0.925–0.993[(C0.999–1S0–0.001)O3]2, and subsequently precipitated low-Mg calcite (LMC), (Ca0.914–0.990 Mg0.006–0.048Fe0–0.003 Na0–0.001Mn0–0.001)0.960–1CO3, and high-Mg calcite (HMC) zones, Mg0.171–0.383Fe0–0.014Na0–0.003Mn0–0.010)0.978–0.992CO3, (Ca0.599–780 deposited in the single grains in the alternate mode. Matrix dolomiteC is a non-stoichiometric (51.0–53.7 mol% CaCO3) dolomite that shows lower contents of Na (b 50–600 ppm), Sr (b50–500 ppm) and S

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(b50–350 ppm), but elevated amounts of Fe and Mn (250 ppm up to 1.58 wt.%), in comparison with the dolomite types A and B. This type of dolomite was not separated from the surrounding LMC and HMC and hence the stable isotope data (− 5.6 to − 5.1‰ for δ18O, − 1.2 to − 1.1‰ for δ13C and +14.2‰ for δ34SCAS) likely reflect mixtures of dolomite and calcite. 4.2.4. Dolomite cement This type of dolomite encounters for about 30–40 vol.% of the estimated total dolomitized rock volume. Its occurrence is restricted to cavities in oolitic to oncolitic grainstones (beds 1–23), where it frequently occurs in the form of medium- to coarse-grained (50–200 μm sized), mostly planar-e shaped and void-filling dolomite rhombs (Fig. 2A). The most prominent feature of the dolomite cement is its fabric-retentive texture and its well-defined zonation pattern, displaying Na-, Sr- and S-rich cores and coarse, partly Fe-rich, rims (Fig. 4D). Dolomite cement is a non-stoichiometric (57.3–60.7 mol% CaCO3) dolomite, which contains moderate to relatively high amounts of Na (50–650 ppm), Sr (b 50–500 ppm), S (b50–1000 ppm) and moderate Fe and Mn (b 50–6800 ppm) contents. The chemical composition of the dolomite cement varies in the range (Ca1.054–1.145Na0.001–0.017 Sr0–0.001)1.061–1.152(Mg0.778–0.900Fe0–0.016Mn0–0.006)0.797–0.900[(C0.997–1 S0–0.003)O3]2. Its isotopic composition ranges from + 1.6 to + 1.8‰ for δ18O and + 1.5 to + 2.0‰ for δ13C, similar to matrix dolomite-A and matrix dolomite-B. 5. Discussion In this section, we revisit the criteria required for the formation of sedimentary dolomite and provide new insights into the physicochemical conditions and processes causing dolomitization of the Oker platform carbonates, based on combined (micro)textural, trace elemental and isotopic data. The importance of BSR during the formation of the Oker dolomite is highlighted and new results about the relationships between (i) the CAS content in dolomite, (ii) its stoichiometry and (iii) the ordering ratio of the dolomite lattice structure at various states of dolomite maturation are presented. Afterwards, we discuss the links between the structure, the composition and the stability of modern and ancient sedimentary dolomites and their fate during burial diagenesis. 5.1. Depositional environments of dolomite formation at Oker Although the abundance of the dolomitized segments in the Upper Jurassic platform carbonate sequence at Oker is apparently higher compared with the limited occurrence of modern dolomites (Morse & MacKenzie, 1990; Holland et al., 1996; Burns et al., 2000), important similarities with modern dolomite-forming environments do exist. For example, the depositional environment, in which the Oker dolomites have been formed, is considered similar to that reported for the present-day carbonate platforms (i.e., dolomite cement) of the Bahama Banks and Florida and for the modern sabkha environments (i.e., dolomite types A to C) of the Coorong Region, Lagoa Vermelha and Abu Dhabi (Tucker & Wright, 1990; Flügel, 2004). In these environments, microbial activity coupled with high Mg/Ca ratios, alkaline pH, high carbonate alkalinity and reduced SO24 − concentrations in the dolomitizing fluids are well-known to facilitate the early diagenetic formation of dolomite in evaporitic, organic-rich sediments dominated by BSR (e.g., Baker & Kastner, 1981; Vasconcelos & McKenzie, 1997; van Lith et al., 2003; Wright & Wacey, 2005; Wacey et al., 2007; Bontognali et al., 2010; Loyd et al., 2012). At Oker, the diversified sedimentary facies and the dominating marine fauna observed across the entire Langenberg section (Fig. 2) indicate that evaporitic and reducing conditions have existed during limestone dolomitization. However, the textural differences between the dolomite cement and the dolomite types A to C (Fig. 4) suggest at least two different dolomitization pathways operating at Oker.

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In the case of the dolomitized oolitic grainstones from Oker (Fig. 2A) the cloudy-center-clear-rim texture (Tucker & Wright, 1990; Budd, 1997), with Na-, Sr- and S-rich cores and partly Fe-rich rims (Fig. 4D and Electronic Appendix B), suggests that the dolomite cement has been formed by the replacement of pre-existing magnesian calcite cement, and subsequent continuous growth of dolomite crystals under reducing conditions (Blake & Peacor, 1985; Budd, 1997; Reinhold, 1998; Warren, 2000; Choquette & Hiatt, 2008; Rameil, 2008). The results from field studies support such a dolomitization pathway in open marine, high-energetic environments such as that of recent ooid shoals, as early void-filling HMC cement is usually dolomitized with good fabric retention, in contrast to fine-grained micrite (Sibley, 1982; Tucker & Wright, 1990). It is likely that the Mg2+ (and Fe2+) ions that are required for the formation of the dolomite cement were supplied by the dissolution of such magnesian CaCO3 precursors, labile Fe(oxyhydr)oxides (Fig. 2A) and/or by reducing, seawater-derived mineralizing fluids. Such physicochemical conditions may have forced a relatively rapid growth of the dolomite crystals (Sibley & Gregg, 1987; Tucker & Wright, 1990), thus promoting the formation of planar and limpid dolomite cement. Such dolomite cements typically form at b35 °C and at shallow (b100 m) burial depths (e.g., Choquette & Hiatt, 2008). In contrast, the characteristics of the massive dolostone beds (Fig. 2B–F) and of the surrounding mudstone support the evidence for a second dolomitization pathway operating in restricted, shallowlagoonal environments. The deposition of micrite, which was subsequently replaced by glauconite group minerals and pyrite, predates the formation of the dolomite types A to C (Fig. 4A–C) in the originally organic-rich sediments, and indicates that the dolomitization fluids were reducing (Baldermann et al., 2012). The strictly marine fauna (Goniolina, Isastrea, and Trigonia sp.) as well as the sedimentary structures (e.g., algae lamination, caliche, tepee structures and birdeyes) observed in the prominent dolostone horizons and in the adjacent mudstone indicate dominating euhaline to hypersaline conditions during the formation of the dolomite and its episodic, tide-related subaerial exposure (Fischer, 1991; Pomoni-Papaioannou, 2008). Evaporitic conditions in shallow lagoons restricted from mixing with seawater could exist temporarily during major sea-level lowstands, which would facilitate the formation of dolomite by high Mg/Ca ratios and high carbonate alkalinity among other factors (e.g., Machel & Mountjoy, 1986; Burns et al., 2000; Warren, 2000; Meister et al., 2013). These observations suggest that the dolomite types A to C were formed by the early diagenetic replacement of micrite under slightly evaporitic and reducing conditions, with major Mg-supply from seawater-derived mineralizing fluids provided by shallow seepage reflux and/or evaporitic tidal pumping (Carballo et al., 1987; Warren, 2000; Qing et al., 2001). Such ancient “penesaline dolomites”, according to Qing et al. (2001), share many textural and isotopic similarities to dolomite crusts known from Holocene tidal flats. Recent tidal-flat dolomites typically form in the presence of sulfate-reducing bacteria and in organic-rich sediments, which contain abundant nucleation sites favorable for a relatively slow growth of the dolomite crystals (Carballo et al., 1987; Tucker & Wright, 1990). Such physicochemical conditions promote the formation of fabricdestructive, dolomite-capped sequences, as observed for the type B of the Oker dolomite (Fig. 2).

5.2. Early versus burial diagenetic origin of the Oker dolomite Most of the sedimentary, non-stoichiometric dolomite is not stable during burial diagenesis and hence is characterized by the reset of its original isotopic and trace elemental signatures (Land, 1980; Morse & Mackenzie, 1990; Malone et al., 1996). Therefore, the key to assess the physicochemical conditions and processes leading to the dolomitization of ancient platform carbonates is to decipher pristine signals from multiple recrystallization events, associated with burial diagenesis

(e.g., Sibley & Gregg, 1987; Price & Sellwood, 1994; Budd, 1997; Reinhold, 1998; Geske et al., 2015). At Oker, the mostly idiotopic dolomite textures (Fig. 4), the small grain sizes of all types of the dolomite (~ 20–50 μm, on average) and the occurrence of perfectly conformable dolostone beds suggest a lowtemperature origin of the dolomite (b40 °C), rather than its formation at elevated temperatures during burial diagenesis (e.g., Sibley & Gregg, 1987; Warren, 2000; Meister et al., 2013). The trace elemental composition (300–400 ppm of Na, ~ 200 ppm of Sr, 360–650 ppm of S and 100–900 ppm of Fe and Mn) of the Oker dolomite supports its early diagenetic origin (Fig. 5A–D), because most dolomites forming at elevated temperatures (N50–60 °C) are usually Mn- and Fe-rich (N1000 ppm), but Na- and Sr-depleted (100–300 ppm) (Sibley & Gregg, 1987; Tucker & Wright, 1990; Price & Sellwood, 1994; Reinhold, 1998; Warren, 2000; Derry, 2010). Importantly, the trace elemental composition of the Oker dolomite is rather similar to that reported for the modern Abu Dhabi dolomites, which form by the early diagenetic replacement of magnesian calcite and aragonite precursors and/or by direct precipitation under evaporitic and partly reducing conditions (Land & Hoops, 1973; Staudt et al., 1994; Budd, 1997; Warren, 2000; Bontognali et al., 2010). The reconstruction of the main depositional environments at Oker supports the concept of dolomite formation and maturation under evaporitic and reducing conditions, and indicates that the boundary conditions for the formation of the Oker dolomite are comparable to that reported from the study of modern sabkha environments. Thus, precipitation and/or extensive recrystallization of the Oker dolomite during burial diagenesis are unlikely. There is no evidence for any modification of the original dolomite signatures, because (i) non-conformable dolomite caps, (ii) sub- to anhedral dolomite cements and (iii) dolomite grains that display significant trace elemental variations between early (non-stoichiometric) dolomite cores and late (more stoichiometric) dolomite generations were not documented in the analyzed samples; features characteristic for the secondary, burial diagenetic dolomite (Sibley & Gregg, 1987; Malone et al., 1996; Reinhold, 1998; Derry, 2010; Geske et al., 2015). 5.3. Timing and chronology of dolomite formation at Oker The δ18O and δ13C isotopic data for the dolomite types A and B, and for the dolomite cement (Fig. 6), plot in the fields of Upper Jurassic inner platform dolomite (Budd, 1997; Rameil, 2008) and modern sabkha dolomite (Warren, 2000; Wacey et al., 2007; Bontognali et al., 2010). The four main types of the Oker limestone plot within the domain of Upper Jurassic inner platform carbonates (Price & Sellwood, 1994; Rameil, 2008). Therefore, the preservation of the pristine isotopic signatures of the platform carbonates from Oker is clearly documented by the range of the δ18O and δ13C isotopes (Fig. 6), with the exception of the samples (i.e. type C dolomite) from the boundary between the mudstone and the upper dolostone. 5.3.1. Type A dolomite It is suggested that matrix dolomite-A represents the earliest generation of dolomite, which is typically formed in spatially restricted microenvironments such as in burrow systems (Fig. 2C), with Mg-supply by shallow seepage reflux and/or evaporitic tidal pumping (Budd, 1997; Gingras et al., 2004; Horbury & Qing, 2004; Rameil, 2008; Meister et al., 2013). The partly Mg- (and Fe2+)-rich organic linings of the burrow, coupled with alkaline pH, high CO2− 3 ion activity and microbial mediation produce conditions ideal for the selective formation of dolomite (Vasconcelos & McKenzie, 1997; van Lith et al., 2003; Gingras et al., 2004; Rameil, 2008). In tidal flats and lagoonal sediments, bacterial colonization of sediment substrates can extend up to several cm's away from the sediment–seawater interface catalyzing a variety of biogeochemical processes, which can induce limestone dolomitization (Purser, 1975; Gingras et al., 2004). The post-depositional formation of such delineated geochemical micro-environments could be responsible

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Fig. 5. Correlations between major and trace elements in the analyzed platform carbonates, revealed by EMP analyses (n = 267). A: Cross-plot of Ca and Mg in a.p.f.u. (atoms per formula unit), including the ideal Ca–Mg substitution line for carbonates (dotted line). B: Enlarged view of A. The four types of the Oker dolomite plot well in the range of (modern, marine) Caexcess dolomites. C: Na versus S cross-plot. D: Na + Sr versus Fe + Mn cross-plot. The samples from the mudstone–upper dolostone boundary (bed 125) show signs of burial diagenetic alteration. See text for explanations.

for the observed irregular spatial distribution of the patchy and/or floating dolomite rhombs, embedded in partly replaced micritic matrix (Figs. 2C and 4A). Although we can only speculate on the absolute δ18O composition of the mineralizing fluids that were involved in the type-A dolomitization, the Δδ18ODol–Cal value (= δ18ODolomite–δ18OCalcite) of ~3‰ (Fig. 6) and the Δδ13CDol–Cal value close to 1‰ imply that micrite precipitated directly from ancient seawater, followed by its instantaneous alteration into matrix dolomite-A without significant deterioration of the isotopic signatures (Sheppard & Schwarz, 1970; Land, 1980; Budd, 1997; Rameil, 2008). Using a δ18O value of − 1‰ (SMOW) for Jurassic seawater (Price & Sellwood, 1994; Reinhold, 1998) and applying the equation of Vasconcelos et al. (2005) the formation temperature of type A dolomite was calculated to be 26 ± 2 °C. 5.3.2. Type C dolomite Matrix dolomite-C could have also been formed by the early diagenetic replacement of pre-existing micrite (Reinhold, 1998). However, the low δ18O values, elevated Fe and Mn contents and higher stoichiometry of this type of dolomite indicate burial diagenetic overprinting, comparing with the other three types of the Oker dolomite (Figs. 5 and 6). The notably lower δ18O values (Fig. 6: − 5.6 to − 5.1‰) of the samples from the boundary between the mudstone and the overlying upper dolostone suggest precipitation of vein-filling calcite spar and concurrent recrystallization of micrite at temperatures up to 75 °C

(Baldermann et al., 2012). Subsequent alteration of type C dolomite, e.g., local de-dolomitization, is evidenced by the above modifications of this type of dolomite (Figs. 4C and 5). 5.3.3. Dolomite cement and type B dolomite The δ18O values of the type B dolomite and of the dolomite cement are relatively high (Fig. 6: + 1.4 to + 2.9‰) and differ significantly from the very low δ18O values (−2.5 up to −16‰) of most burial dolomites (e.g., Reinhold, 1998; Derry, 2010). The Δδ18ODol–Cal values of ~3 up to ~5‰ support an early diagenetic origin of these two types of the Oker dolomite and indicate that they were likely formed on the extent of pre-existing HMC and LMC through interaction with slightly evaporitic seawater-derived mineralizing fluids (Warren, 2000; Qing et al., 2001; Rameil, 2008). Despite the apparently different textures of the dolomite cement and of the type B dolomite (Fig. 4B and D), their almost identical isotopic and trace elemental composition (Figs. 5 and 6) suggests that they were formed simultaneously and/or in rapid succession from similar mineralizing fluids at low to moderate temperatures. This suggestion is supported by the estimated paleo-temperatures from 22 ± 3 °C for the pre-existing micrite and from 37 ± 2 °C for these two types of the Oker dolomite (Baldermann et al. 2012). Moreover, these formation temperatures are in the range of the estimated paleotemperatures for the Upper Jurassic platform carbonates from the central Jura Mountains (NW Switzerland) and Ammonitico Rosso from

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Fig. 7. δ34SCAS data (n = 8). The dolomite types A and B reveal pristine-marine δ34SCAS values, whereas the matrix dolomite-C from the mudstone–dolostone boundary (bed 125) shows signs of diagenetic overprinting. See text for explanations. Fig. 6. Cross-plot of the δ18O and δ13C isotopes. The Oker limestone plots within the field of Upper Jurassic inner platform carbonates. Matrix dolomite-A, matrix dolomite-B and dolomite cement from Oker plot in the range of Upper Jurassic sabkha dolomites and recent Abu Dhabi sabkha dolomites. The δ18ODol–Cc values (= δ18ODolomite–δ18OCalcite; Budd, 1997) of ~3 up to 5‰ indicate that micrite precipitated from Upper Jurassic seawater, subsequently transformed into dolomite through interaction with slightly evaporatic, seawater-derived interstitial solutions. The samples from the boundary between the mudstoneand the overlaying upper dolostone show significantly lower δ18O values compared to the pristine platform carbonates from Oker, which suggests local alteration during burial and/or meteoric diagenesis.

Cala Fornells (Mallorca) (Price & Sellwood, 1994; Rameil, 2008). Moderate formation temperatures (b40 °C) have been reported for the modern sabkha dolomites from Abu Dhabi and from the Coorong Region (Warren, 2000; Bontognali et al., 2010).

5.4. The role of BSR on the formation of dolomite The relationship between the secular variations in the chemical composition of seawater, BSR and the occurrence of volumetrically significant dolostone deposits throughout the geological record has been intensively discussed (Burns et al., 2000; Warren, 2000). It has been proposed that a greater extent of sulfate reduction in calcareous sediments should increase the rate of limestone dolomitization (e.g., Claypool et al., 1980; Wright, 1999; Burns et al., 2000). Recently, the δ34SCAS signatures of marine carbonates have been used for reconstructing the sulfate–sulfur isotopic composition (δ34SSO4) of ambient seawater (Kampschulte & Strauss, 2004; Paytan et al., 2011). Kampschulte and Strauss (2004) determined the δ34SSO4 composition of Kimmeridgian seawater to be ~+17‰. In the case of the Oker dolomites, the shift towards positive δ34SCAS values, from + 17.9‰ for the early matrix dolomite-A up to + 19.7‰ for the evolved matrix dolomite-B, supports the evidence for BSR accompanying the precipitation and the growth of dolomite (Fig. 7). In accordance, near-seawater δ34SCAS values (+ 19.0 to + 21.8‰, V-CDT, compared to an estimated Miocene seawater δ34SSO4 value of ~ 22‰, V-CDT) or positive δ34SCAS excursions (19.0 up to 36.9‰, V-CDT) have been reported for the partly dolomitized carbonate concretions from the Miocene Monterey Formation, which have been formed within the zone of sulfate reduction and/ or during methanogenesis (Loyd et al., 2012). Thus, the positive shift of ~2‰ for the δ34SCAS values of the Oker dolomite (Fig. 7) follows a change in the δ34SSO4 isotopic composition of the dolomitizing fluids and reflects the involvement of BSR processes (e.g., Kampschulte & Strauss, 2004; Loyd et al., 2012; Gomes & Hurtgen, 2015). At the time of dolomiconcentrations, which tization BSR is reducing the pore water SO2− 4 is documented by the decreasing abundance of CAS in precipitating

dolomite (Fig. 8A), and finally results in the syngenetic formation of euhedral pyrite (Fig. 4A). The experimental work of Malone et al. (1996) suggests that the recrystallization of early (poorly ordered and non-stoichiometric) dolomite cores into later (more stoichiometric) dolomite generations leads to a reduction of the trace element content in dolomite, and the removal of CAS from the dolomite crystal lattice could affect the ordering ratio of the dolomite. At Oker, the low Na, Sr and CAS contents (550 mg/kg, as determined by EMP analyses) reported for the altered samples from the boundary between the mudstone and the overlaying upper dolostone suggest a diagenetic modification of the type C dedolomite (Fig. 5C and D), which is also supported by their notably lower δ18O values (Fig. 6). In contrast, Loyd et al. (2012) and Gomes and Hurtgen (2015) have demonstrated that the CAS content of marine carbonate minerals, which form near the onset of the zone of sulfate reduction, is in the range of the sulfate concentration of the ambient seawater. Thus, the 1400 mg/kg and 1100–1950 mg/kg of CAS measured for the type A and type B of the Oker dolomite, respectively, suggest insignificant alteration of the pristine CAS contents, considering the estimated ~15–20 mM sulfate concentration of Upper Jurassic seawater (e.g., Algeo et al., 2014). Certainly, the CAS content of marine carbonates is strongly dependent on parameters such as the precipitation rate, mineralogy and the amount and type of those minerals in sediments (Busenberg & Plummer, 1986; Loyd et al., 2012; Gomes & Hurtgen, 2015). The experimental work of Busenberg and Plummer (1985) has shown that incorporation of Na+ and SO2− 4 ions in precipitating calcite is largely dependent on the calcite precipitation rate. Such a covariance contents is also evident for the Oker dolomite of the Na+ and SO2− 4 (Figs. 5C and 8A), which indicates that both ions have been incorporated during dolomite precipitation and growth. Accordingly, the reported CAS contents are suggested to reflect pristine signatures of dolomite formation with no recognizable later modification. These observations are one of the key aspects to the understanding of dolomitization of the Upper Jurassic platform carbonates at Oker. Sulfate is a well-known inhibitor for the homogeneous nucleation and concentracontinuous growth of dolomite and thus decreasing SO2− 4 tions in the mineralizing fluids should facilitate the formation of dolomite (Baker & Kastner, 1981; Vasconcelos et al., 1995; Vasconcelos & McKenzie, 1997; Warthmann et al., 2005; Deng et al., 2010; Loyd et al., 2012). In order to model the effect of BSR on the speciation and saturation state of seawater in respect to dolomite, the PHREEQC software together with its LLNL database was used (Parkhurst & Appelo, 1999). The results of the speciation calculations reveal that a decrease 2+ of the SO2− 4 concentration in seawater causes an increase of free Mg and Ca2+ ions originally forming MgSO04 and CaSO04 aqueous species,

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Fig. 8. A: Cross-plot showing the sulfur content of the Oker dolomites as determined by EMP analyses and CAS extractions. The plot indicates that sulfur in dolomite is present only in the form of CAS. B: Effect of ongoing sulfate reduction on the molar ratio of free aqueous Mg2+/Ca2+ ions (black curve) and on the resultant saturation index (SI) of dolomite with respect to seawater (gray curve), calculated with the PHREEQ-C software and its implemented LLNL-database.

and leads to a further increase of dolomite supersaturation (Fig. 8B). This reaction chain can be viewed as the abiotic part of the microbial model, which has been suggested to be applicable to the precipitation of dolomite in modern sabkha environments (Vasconcelos & McKenzie, 1997; van Lith et al., 2003; Wright & Wacey, 2005; Wacey et al., 2007; Bontognali et al., 2010). However, the experiments of Sánchez-Román et al. (2009) and Deng et al. (2010) on microbially mediated dolomite formation in the presence and absence of sulfate have shown that bacterial activity mediates the precipitation of nearly stoichiometric, but almost completely disordered (proto)dolomite. Their experimental results suggest that, besides BSR, other co-existing factors could be important for the formation of dolomite in bacteria-dominated environments, such as (i) the presence of EPS secreted by bacteria, (ii) the pH and the carbonate alkalinity of the dolomitizing fluids, which directly affect the supersaturation degree in respect to dolomite and (iii) appropriate templates for the nucleation and growth of dolomite (Lippmann, 1973; McKenzie, 1981; van Lith et al., 2003; Bontognali et al., 2010; Loyd et al., 2012; Meister et al., 2013). The results of the recent experimental work of Bontognali et al. (2014) support the evidence for EPS being a key factor in the microbially influenced formation of Ca–Mg-carbonates at low temperatures, and further indicate that an increase in pH and alkalinity, and sulfate removal due to microbial respiration appear less important than suggested in earlier studies. The reaction mechanism proposed includes aquocomplex formation of Ca2+ and Mg2+ by EPS at considerable concentrations, which subsequently leads to the formation of spatially restricted

micro-environments within the microbially produced biofilms that are characterized by elevated Mg/Ca ratios and a high degree of dolomite supersaturation with respect to the bulk solution. It has been suggested that organic polymers can increase the dehydration of Mg2+ aquocomplexes on the crystal surface and thus result in the preferential incorporation of Mg2+ ions into the precipitating (proto)dolomite (Bontognali et al., 2014). Although the active role of EPS in the mineralization process of the Oker dolomite cannot be inferred from the obtained results, the involvement of sulfate-reducing bacteria during dolomite formation and its subsequent growth is clearly indicated by the decreasing CAS contents and the positive δ34SCAS excursions as the dolomite matures (Figs. 7 and 9). The interplay of the above biotic and abiotic factors seems to control the underlying dolomitization pathways in low-temperature environments and subsequently predefines the dolomite textures and the abundance of dolomite in modern and ancient sediments. 5.5. Links between the CAS content and the composition, structure and stability of modern and ancient sedimentary dolomites In modern dolomite-forming environments, dolomitization progresses in originally organic-rich sediments either by direct precipitation or by the early diagenetic replacement of magnesian calcite and/ or aragonite precursors at low to moderate temperatures (b40 °C) and in the presence of sulfate-reducing bacteria, with the essential aqueous Mg being supplied by shallow seepage reflux and/or evaporitic tidal

Fig. 9. Degree of cation order in sedimentary dolomite (Dol) plotted against (A) the CAS content and (B) dolomite stoichiometry. The CAS composition and the estimated ordering ratio of the dolomite from the Miocene Monterey Formation are adapted from Loyd et al. (2012).

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pumping (Vasconcelos & McKenzie, 1997; Warren, 2000; Warthmann et al., 2005; Wacey et al., 2007; Bontognali et al., 2010). Although apparently similar physicochemical conditions for the formation of sedimentary dolomite are verifiable at Oker and in modern sabkha environments, some important differences do exist concerning the mineralogical composition and the structure of the precipitating dolomite. For instance, most modern (proto)dolomites are non-stoichiometric (Ca excess) and characterized by a disordered structure, whereas ancient dolomites that have underwent aging processes, i.e. multiple recrystallization during burial diagenesis, are usually near-stoichiometric and show at least a partial ordering (ordering ratio N0.3) of the dolomite lattice structure (e.g., Malone et al., 1996; Reinhold, 1998; Burns et al., 2000; Warren, 2000). At Oker, overprinting of the original signatures of the dolomites during the burial period could be almost excluded by the range of obtained trace elemental and isotopic data (Figs. 5–7), except for the type C dedolomite. These findings suggest that the textural and structural variations between the immature type A dolomite and the evolved type B dolomite and the dolomite cement (Fig. 4) can be directly interpreted in terms of different states of dolomite maturity linked to its formation at relatively low to moderate (~ 26–37 °C) temperatures. For the first time, we show here that a linear anti-correlation exists between increasing ordering ratio of the dolomite lattice structure, from 0.34 to 0.50, and decreasing CAS content in dolomite, from 1950 to 1100 ppm (Fig. 9A). This relationship is valid for the Miocene Monterey dolomites as well and thus it can be used as a measure for dolomite maturity (Fig. 9A and B), according to the expression:   degree of cation orderðDolÞ ¼ −0:018  CASðDolÞ þ 68:3 R2 ¼ 0:98 : In marine-diagenetic environments, the CAS content and subsequently the ordering ratio of the dolomite precipitating within the zone of sulfate reduction are likely to be controlled both by the sulfate concentration of the dolomitizing fluids and the sulfate reduction rate in marine sediments (Loyd et al., 2012; Gomes & Hurtgen, 2015). Considering the relatively high sulfate concentrations (34–100 up to 590 mM) in most modern dolomite-forming environments (Warren, 2000; Deng et al., 2010) it is suggested that the precipitating (proto)dolomites contain high CAS contents (N 2000 mg/kg), which would explain their poorly ordered or almost completely disordered (ordering ratio b0.3) structure (Fig. 9A). Yet, there exist no studies in which the links between the CAS content, the structure and the stability of modern (proto)dolomite have been systematically investigated. It is well-known that the thermodynamic properties of sedimentary dolomites largely depend on the degree of ordering of the dolomite lattice structure and on its stoichiometry. The Gibbs free energy (ΔGf) of disordered (proto)dolomite is more than 1.3 kcal/mol lower than that of ordered dolomite (at 25 °C and 1 atm), which is equivalent to a difference of about one and a half orders of magnitude in solubility constants for both types of the sedimentary dolomite (Warren, 2000). Consequently, disordered (proto)dolomite is the most soluble and less stable form of dolomite in marine-diagenetic environments and it tends to recrystallize over time and during burial diagenesis into structurally and compositionally more stable forms of dolomite, thereby resetting its original isotopic and trace elemental signatures (e.g., Land, 1980; Warren, 2000; Geske et al., 2015). It is suggested that during the periods of Earth's history (e.g., in the Upper Jurassic) where the sulfate concentration of the seawater was low and/or the rate of BSR was high (e.g., Kampschulte & Strauss, 2004; Algeo et al., 2014), the precipitating dolomite should be relatively poor in CAS and thus more stable during burial diagenesis, as evidenced by the preservation of the original (micro)textural and (isotope)geochemical signatures of the Oker dolomites (Figs. 5–7). The relationships between the sulfate concentration in the dolomitizing fluids, the rate of BSR in sediments and subsequently the CAS content of the precipitating dolomites are suggested

to play a key role in controlling the composition and structure of modern and ancient sedimentary dolomites and hence their stability during burial diagenesis. 6. Summary and conclusions The Langenberg section at Oker comprises the most complete succession of Upper Jurassic rocks in Central Germany. The ~150 m thick sedimentary sequence is dominated by partly dolomitized, oolitic grainstones and Exogyra boundstones, deposited along an open marine ooid shoal and barrier-reef facies, which delineated a shallow-lagoonal, inner carbonate platform setting that is characterized by alternate argillite, mudstone and dolostone. Based on detailed field work and unique dolomite textures, three types of early diagenetic dolomite (i.e., matrix dolomite-A, matrix dolomite-B and dolomite cement) and one type of dedolomite (matrix dolomite-C) were distinguished. Our combined (micro)textural and (isotope)geochemical data set of the platform carbonates from Oker provide a strong foundation for the following physicochemical conditions and processes leading to dolomite formation at low temperatures. (i) The limestone (−3.4 to +0.1‰ for δ18O and −4.8 to +2.6‰ for δ13C) and the dolostone (+ 1.4 to + 2.9‰ for δ18O and 0.1 to + 2.0‰ for δ13C) from Oker are clearly distinguished by their stable oxygen and carbon isotopic composition. Fractionation of δ18O isotopes between both calcite and dolomite of about 3 up to 5‰ suggests that the micritic calcite precipitated directly from ancient seawater (except for the carbonates from the mudstone–upper dolostone boundary), followed by its early diagenetic transformation into dolomite. (ii) The nearly identical δ18O and δ13C isotopic composition (+1.4 to + 2.9‰ for δ18O and − 0.1 to + 2.0‰ for δ13C) of the dolomite cement and of the dolomite types A and B suggests that the Oker dolomites could have been formed from similar dolomitization fluids at moderate temperatures between 26 °C and 37 °C. Pristine marine to slightly evaporitic and reducing seawaterderived interstitial solutions are suggested as the dolomitization fluids, which were provided by seepage reflux and/or evaporitic tidal pumping. (iii) The sulfur isotopic composition of CAS highlights the incorporation of sulfate–sulfur during dolomite precipitation from ambient Kimmeridgian seawater, where elevated δ34SCAS values from +17.9 to + 19.7‰ suggest BSR during dolomite formation. In the case of the Oker dolomites, a higher degree of cation order in the dolomite lattice structure anti-correlates linearly with decreasing sulfate–sulfur contents. This relation can be directly used as a measure for the ordering ratio of the dolomite lattice structure. (iv) Spatial trace elemental distribution on the micron-scale verifies the concept of precipitation and maturation of dolomite during ongoing BSR, as zoned dolomite crystals display Na- (300– 400 ppm), Sr- (~200 ppm) and S-rich (350–650 ppm) cores and partly Fe- and Mn-rich (100–900 ppm) rims. Such trace element patterns indicate that matrix dolomite-A could have precipitated first, followed by continuous growth of matrix dolomite-B and dolomite cement under reducing conditions. (v) An impact of meteoric solutions during limestone dolomitization can be almost excluded by the range of oxygen isotope data. Only the samples from the mudstone/upper dolostone boundary, which comprise secondary calcite spar and matrix dolomite-C, show signs of localized meteoric or burial diagenetic overprinting, as documented by lower δ18O values from − 5.6 to −5.1‰. In summary, the proposed conceptual model for limestone dolomitization at Oker includes direct precipitation of magnesian CaCO3

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precursors from ambient Upper Jurassic seawater, which are subsequently transformed into dolomite at low temperatures, with Mg supply from pristine marine to slightly evaporated seawater-derived mineralizing fluids. We conclude that BSR is an important process during dolomitization, as the CAS content in dolomite decreases from poorly ordered to well-ordered dolomite. The anti-correlation between the CAS content in dolomite and its ordering ratio might explain why most modern marine dolomites that form in low-temperature environments are metastable and show only partial ordering of the lattice structure. We further conclude that ongoing reduction of pore water sulfate is responsible for the early diagenetic mobilization of Fe2+ ions, which is a precondition for the often reported subsequent precipitation of pyrite, glauconite group minerals and Fe-rich later generations of dolomite at reducing conditions. Supplementary data to this article can be found online at http://dx. doi.org/10.1016/j.chemgeo.2015.07.020.

Acknowledgments The funding by the NAWI Graz is highly appreciated. We are grateful to D.R. Hilton, T. Bontognali and an anonymous reviewer for their constructive and insightful reviews.

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