dolomitization of lower cretaceous peritidal … · 2015. 12. 14. · using the sibley and gregg...

Journal of Sedimentary Research, 2014, v. 84, 552–566

Research Article

DOI: http://dx.doi.org/10.2110/jsr.2014.45

DOLOMITIZATION OF LOWER CRETACEOUS PERITIDAL CARBONATES BY MODIFIED SEAWATER:CONSTRAINTS FROM CLUMPED ISOTOPIC PALEOTHERMOMETRY, ELEMENTAL CHEMISTRY, AND

STRONTIUM ISOTOPES

CLAIRE M. SENA,*1 CEDRIC M. JOHN,1 ANNE-LISE JOURDAN,{1 VEERLE VANDEGINSTE,1 AND CHRISTINA MANNING2

1Department of Earth Science and Engineering, Royal School of Mines, Imperial College London, London SW7 2BP, U.K.2Department of Earth Sciences, Queens Building, Royal Holloway University of London, Egham TW20 0EX, U.K.

e-mail: [email protected]

ABSTRACT: This study investigates one of the few examples of dolomitization of Lower Cretaceous shallow-water limestonesfrom the southern Tethys carbonate platform from outcrops on the Haushi-Huqf High in central-east Oman. Mud-dominatedperitidal carbonates rich in microbial mats are replaced by fine crystalline dolomite along at least 60 km in the lower 10 m of theJurf Formation. Two, meter-thick beds in the overlying Qishn Formation are also dolomitized over lateral distances of one to twokilometers. The stratabound geometry and petrographic relations of the dolomite with all other diagenetic phases indicate that thedolomite precipitated early. The presence of only rare anhydrite relicts suggests that seawater was below gypsum saturationduring most of the dolomitization, supporting the hypothesis that Cretaceous slightly evaporated water (d18Ofluid values rangefrom 2.8% to 3.5% SMOW) can affect pervasive dolomitization. Despite being composed of peritidal facies with featuressuggesting high salinity, most of the carbonate succession was not dolomitized, suggesting that the presence of microbial matsexerted a major control on the distribution of dolomite, and that salinity is only a minor control on dolomitization. Clumped-isotope results indicate that the early formed dolomite re-equilibrated with fluids at 44 ± 3uC in a shallow-burial setting. Theelevated iron content (on average 6011 ppm) and 87Sr/86Sr (on average 0.70782) with respect to Cretaceous seawater suggests thatthe burial fluids interacted with Permian clay and feldspars of the Gharif Formation that underlies the Cretaceous carbonates. Thevertical gradients of radiogenic Sr and the occurrence of small volumes of coarse crystalline dolomite in stylolites and fracturessuggests that burial fluids were driven vertically and laterally by differential compaction, probably during maximum burial in theLate Cretaceous.

INTRODUCTION

Dolomitization of shallow-water limestone in near-surface and shallowburial settings, although common in the stratigraphic record, isparticularly controversial given the failure to precipitate dolomite fromsupersaturated solutions in abiotic laboratory experiments under near-surface conditions (Lippman 1973; Land 1998). Fundamentally, anymodel for the formation of large dolomite bodies must satisfy two basiccriteria: 1) thermodynamic and kinetic conditions must be favorable todolomitization, and 2) there must be a fluid-flow mechanism by whichreactants and products can be transported to and from the site ofdolomitization (Land 1985; Machel and Mountjoy 1986; Hardie 1987).

Several conceptual hydrogeological models for the formation of earlydolomite were proposed based on observations of ongoing dolomiteformation in modern hypersaline sabkhas in the Persian Gulf (McKenzieet al. 1980) and in Cenozoic carbonate islands (see review in Budd 1997).More recently, it has been found that microbial activity and the presenceof exopolymeric substances may control the formation of early dolomiteby breaking the kinetic barriers leading to dolomite precipitation(Vasconcelos and McKenzie 1997; Bontognali et al. 2010).

Understanding the origin of replacement dolomite in ancient succes-sions is key to improve knowledge on dolomitization processes, and torefine predictive models on dolomite distribution. The study of earlydolomite in epicontinental seas is particularly interesting, given that thetopographic and hydrodynamic settings of epeiric platforms may nothave been analogous to the hydrologic system of modern carbonatesystems (Allison and Wells 2006).

This paper investigates outcrops in the Haushi-Huqf High, central-eastOman (Fig. 1), one of the few examples of dolomitization of LowerCretaceous shallow-water limestone deposited in the southern Tethyscarbonate platform. The Barremian–Aptian Jurf and Qishn formationshave experienced little burial (200 6 50 meters, Immenhauser et al.2004); exposures are excellent. Field relations show that about two-thirdsof the total thickness of the Jurf Formation is dolomitized, dolomite unitsbeing , 10 m thick and possibly can be traced over 60 km laterallyaccording to stratigraphic data (Immenhauser et al. 2004) (Fig. 2). Thepresence of a laterally extensive dolomite body in the Jurf Formation isintriguing, given the scarce and small volumes of dolomite reported foroutcrops of time-equivalent units (Kharaib and Shu’aiba formations) tothe north of the Haushi-Huqf High in Jebel Madar (van Buchem et al.2002; Le Bec 2003) and in the subsurface to the west in the Oman Interiorbasins (Droste 2010).

The goals of this paper are (1) to document the distribution of dolomitetextures in the Jurf and Qishn formations cropping out in Wadi Baw, and

* Present Address: Shell U.K. Ltd, 40 Bank Street, London E14 5NR, U.K.{ Present Address: Bloomsbury Environment Isotope Facility, Department of

Earth Sciences, University College London, London WC1E 6BT, U.K.

Published Online: July 2014

Copyright E 2014, SEPM (Society for Sedimentary Geology) 1527-1404/14/084-552/$03.00

FIG. 1.—A) Map of the Arabian Peninsula,red rectangle indicates the location of theHaushi-Huqf outcrops. B) Simplified geologicalmap from the Haushi-Huqf area adapted fromPlatel et al. (1992) showing the location of WadiBaw (red rectangle).

FIG. 2.—Stratigraphic column of the Haushi-Huqf area showing the dolomite distribution inthe Early Cretaceous carbonates. Timescalebased on Gradstein et al. (2004). Field photo-graphs show the distribution of dolomite withinthe sequence stratigraphic context (names ofsequences as in Sena and John 2013).

DOLOMITIZATION OF PERITIDAL LIMESTONES 553J S R

to investigate how the dolomite textures relate to the depositional fabricof the Barremian–Aptian shallow-water carbonates, (2) use a combina-tion of petrographic and geochemical techniques, including clumpedisotope paleothermometry, to investigate the origin of diagenetic fluidsand the sequence of events leading to dolomitization, and (3) to discussthe implications of findings for dolomitization processes in general andfor the potential occurrence of dolomite elsewhere in the LowerCretaceous Arabian platform.

STRUCTURAL AND STRATIGRAPHIC SETTING

The Haushi-Huqf High is a major anticlinal structure elongated in theNNE–SSW direction that is ca. 40 km wide and extends for ca. 180 kmfrom near Duqm to the Wahiba Sands (Fig. 1) (Ries and Shackleton1990). The Lower Cretaceous sedimentary basin in Oman was subjectedto transtensional and transpressional tectonic regimes related to themovement of the Indian plate during the Mesozoic (first Alpine phase)and the collision between the Arabian and Iranian plates during theCenozoic (second Alpine phase) (Filbrandt et al. 2006). The centraleastern Oman margin was subjected to left-lateral transform faulting andhad a relatively quiet tectonic history compared to the Oman Mountainsin the north, which evolved in a compressional deformation zone.

Rocks exposed in the Haushi-Huqf area are composed of Ordovician toCretaceous sequences with numerous unconformities that onlap foldedVendian–Cambrian Huqf Group strata (Ries and Shackleton 1990; Platelet al. 1992). The Huqf area has experienced low rates of subsidence andintermittent uplift since the Cambrian (Visser 1991). By contrast, age-equivalent sedimentary rocks in the Oman Interior basins underwentsubstantial subsidence, and Precambrian deposits presently lie at a depthof , 10 km (Loosveld et al. 1996). The Lower Cretaceous rock recordconsists of a westward-thickening wedge, from the Haushi-Huqf High tothe south of Oman and the Ghaba Salt basins (Loosveld et al. 1996). Inthe Haushi-Huqf area, the Lower Cretaceous rocks lie unconformablyover Permian sandstones of the Gharif Formation and are unconform-ably overlain by Albian shales from the Nahr Umr Formation (Fig. 2).

Shallow-water and restricted oceanic circulation around the Haushi-Huqf topographic high gave rise to a long-lasting peritidal environment.The Jurf Formation forms one sequence deposited during the earlyBarremian, and is composed of stacked tidal-flat facies (Immenhauser etal. 2004). The Qishn Formation was deposited in peritidal and subtidalenvironments during late Barremian to early Aptian, and records moreopen marine conditions than the Jurf Formation. Subaerial unconformi-ties and sedimentation patterns of the Qishn Formation reflect relativechanges in sea level resulting in repeated subaerial exposure of theplatform (Sena and John 2013). The dominance of miliolids and oysters,and the exclusion of rudists in the peritidal environment, combined withthe occurrence of normal marine biota (echinoids, corals, green algae)towards the subtidal platform, suggest a platform-wide salinity gradient,with increasing salinities toward the platform interior (Immenhauser et al.2004; Sena and John 2013).

METHODS

Five stratigraphic sections (informally named I, J, A, B, G, in order ofincreasing distance from the Haushi-Huqf High) were sampled on a bed-by-bed basis along a NW–SE 16.5-km-long transect in Wadi Baw, south ofthe Haushi-Huqf High (Fig. 1). Sections I, J, A, and B could be confidentlycorrelated in the field by tracing beds laterally, whereas there is ca 2.5 metervertical uncertainty on the stratigraphic correlation with section G.

A total of 244 thin sections were examined under transmitted light andusing a CITL Cathodoluminescence Mk5-2 stage mounted on a NikonEclipse 50i microscope with an attached Nikon DS-Fi1c digital camera.Operating conditions were about 270 mA and 14 kV. Half of each thin

section was stained with alzarin red S and potassium ferricyanide todistinguish calcite from dolomite and their ferroan equivalents following aprocedure modified from Dickson (1966). Dolomite texture was describedusing the Sibley and Gregg (1987) classification scheme that employs twomain categories: 1) crystal size distribution (unimodal or polymodal), and2) crystal boundary shape (nonplanar, planar-s, and planar-e). Furtherdescription of dolomite includes the type of replacement of allochems (non-mimetic or mimetic), types of porosity (following the classification ofChoquette and Pray 1970), and the nature of the minerals filling porosity.The white-card and blue-light fluorescence techniques (Dravis 1991) wereused to resolve depositional textures in dolomites.

In total, 244 bulk powders (1–2 g) were prepared for X-ray diffraction(XRD) by crushing the samples using an agate mortar and pestle. XRDanalysis was carried out with a Philips PW 1830 diffractometer systemusing CuKa radiation at 45 kV and 40 mA, a PW 1820 goniometer, and agraphite monochromator. The scanning range was 2.5u to 70u 2h with astep of 0.01u 2h and acquisition time of 2 seconds per step. Differentmineral phases were quantified using peak intensity ratios, and CaCO3

mol% of dolomite was determined using the equation of Lumsden (1979)and an internal halite standard. The instrumental error of XRD forCaCO3 mol% is 6 0.33 mol %.

Oxygen and carbon isotopes and clumped isotopes were measured inthe Qatar Stable Isotope Laboratory at Imperial College London. d18Oand d13C values for 100 samples were measured from bulk powders(100 mg) composed of 10 to 100% dolomite mixed in the carbonatefraction (as determined by XRD). Samples were reacted with 105%orthophosphoric acid at 70uC in a Kiel IV carbonate device, and theresulting gas was analyzed on a Thermo Finnigan MAT 253 massspectrometer. Data are reported using the standard per mil notation (%)in the VPDB (Vienna Pee Dee Belemnite) reference frame. Datacorrections for instrumental drift were based on multiple runs of theNBS 19 international standard and an internal laboratory standard(Imperial College Carrara marble, ICM). Precision (1 standard errorS.E.) for carbonate standards is better than 0.02% for d13C and 0.04%for d18O. The oxygen isotope composition of 100% dolomite samples wascorrected for acid fractionation using the equation given by Rosenbaumand Sheppard (1986). No correction was made for dolomitic samplescontaining more than 5% calcite in the carbonate fraction. No othercorrections have been applied to d18O given the uncertainty on dolomite–water, calcite–water fractionations (Land 1980) and on fractionationrelated to dolomite stoichiometry (Vahrenkamp and Swart 1990).

Three samples of pure dolomite (based on XRD data) from the Jurf(samples I4 and G5) and Qishn formations (sample E36) were selected forclumped-isotope analysis. Around 8 mg of dolomite were reacted for 1h ina phosphoric acid bath held at 90uC, and the liberated CO2 gas was purifiedby passage through a conventional vacuum line with multiple cryogenictraps and a Porapak-Q trap held at 235uC (following Dennis and Schrag2010). The clean CO2 was analyzed using a Thermo Finnigan MAT 253mass spectrometer capable of measuring molecular masses in the range 44–49 simultaneously. This setup allows the measurement of conventionalisotopes (d18O and d13C) and clumped isotopes (D47) on the same aliquot ofsample. Values of D47 (i.e., the abundance of mass 47 isotopologues13C18O16O over a stochastic distribution in a given sample) were calculatedfrom measured ion intensity ratios and corrected for nonlinearity using theheated-gas-line method described in Huntington et al. (2009). All valueswere corrected for temperature-dependent acid fractionation by adding afactor of +0.081% (Bristow et al. 2011), and are reported using the per milnotation (%) in the absolute reference frame (ARF, Dennis et al. 2011).Masses 48 and 49 were monitored for evidence of sample contamination.Between 3–4 aliquots of the same sample were measured to improve precision,which is around 0.004–0.005% (1 standard error, S.E.). The D47 ARF valueswere converted to temperature using the calibration of Ghosh et al. (2006)established using calcite and aragonite precipitates.

554 C.M. SENA ET AL. J S R

TABLE 1.— Facies classification, sedimentological and depositional environment interpretation. The carbonate textures are defined using the Dunham(1962) classification.

FaciesCode Name Texture Components Structures Depositional Environment

FA1: Low-diversity fauna M-W

1a Laminated mudstone towackestone

M-W skeletal debris (1), porifera (1), blackpebbles

mm-scale sub-parallel wavy laminae,root structures, birds-eye structures,polygonal desiccation cracks, tepeestructures, brecciation, dissolutioncavities, geopetals

Supratidal tidal flat.Exposure index 90–100.

1b Bioturbated wackstonewith packstone–grainstonepatches

W (P-G) skeletal debris (2–3), peloids (1–2),pyrite (1), porifera (2), dasycladacean(1–2), miliolids (1–2)

bioturbation, root structures, geopetalvadose silt, discontinuous packstone–grainstone beds with reworked biota,birds-eye structures, desiccation cracks

Storm influenced intertidaltidal flat. Exposure index55–100.

1c Bioturbated low-diversitywackestone

W miliolids (3), other foraminifera(1–2), skeletal debris (2),porifera (1–2)

bioturbation, desiccation cracks Intertidal tidal flat withrestricted conditions.Exposure index 55–100.

1d Mixed siliciclastic carbonate W quartz (3), moulds of bioclasts (1) horizontally aligned or patchydistribution of quartz grains

Supratidal to intertidal tidalflat.

1e Microbial laminites M microbial laminites (3) desiccation cracks, wavy crinklylaminations (Fig. 3A, B), fenestralstructures, anhydrite relicts (Fig. 3C)

Supratidal to intertidal tidalflat.

FA2: Low-diversity fauna F-G

2a Cross-laminated grainstone G peloids (3), abraded skeletaldebris (3), oysters (2), pyrite (1)

dm-scale low angle bi-directional troughcross-bedding, desiccation cracks

High energy intertidalplatform.

2b Oyster floatstone withpackstone matrix

F oysters (3), peloids (2–3), skeletaldebris (2), other foraminifera (1–2),miliolids (1–2), dasycladacean (1–2)

alignment of shells, erosive base,bioturbation

High energy intertidalplatform.

FA3: High-diversity fauna W-G

3a Bioturbated high diversitywackestone to packstone

W-P echinoderms (1), gastropods (1),miliolids (1), other foraminifera (1–2),reworked coral heads (1),dasycladacean (2–3), skeletal debris(2), peloids (2), orbitolinids (1–2)

abundant bioturbation Medium-energy shallowsubtidal platform.

3b Coral biostrome G peloids (3), oysters (2), in situ coralheads (2), miliolids (2), skeletaldebris (2), rudists (1)

Coral biostrome in shallowsubtidal platform.

3c Bioturbated high diversitywackestone

W skeletal debris (3), echinoderms (2),gastropods (2), dasycladacean (2),porifera (1)

abundant bioturbation Low energy protectedplatform.

FA4: Rudist-dominated F-R

4a Bioturbated floatstone F rudists (2–3), dasycladacean (2–3),echinoderms (2), other foraminifera(1–2), peloids (1–2), porifera (1),miliolid (1), skeletal debris (1)

bioturbation, dm-scale tabular beds Medium energy subtidalplatform.

4b Cross-laminested rudstone R abraded skeletall debris (3),peloids (3), rudists (2), oyster (2),other foraminifera (1–2),dasycladacean (1)

m-scale planar cross-bedding, swellingerosive base and eroded tops, poorlysorted sediments, load structures

High energy subtidalplatform.

4c Bioturbated rudstone R skeletal debris (3), peloids (2–3),dasycladacean (1–2), reworked coral (1)

bioturbation, dm-scale tabular beds High energy subtidalplatform.

4d Rudist biostrome F in situ rudists (3), skeletal debris (3),echinoderms (2), peloids (1–2),oysters (1–2)

Rudist biostrome in subtidalplatform.

FA5: Orbitolinid-dominated P-R

5 Bioturbated packstone-rudstone with orbitolinids

P-R orbitolinids (3), in situ coral heads (2),skeletal debris (2), peloids (2),dasycladacean (1)

bioturbation High energy subtidalforeshoal platformpermanently exposed tocurrent and waves.

FA6: Argillaceous W-P (G)

6a High-diversity fauna grainstone G skeletal debris (3), peloids (3),echinoderms (2–3), benthicforaminifera (2–3), Chofatelladecipiens (2), porifera (2)

cm-scale wavy beds with erosional bases Storm influenced deepsubtidal platform abovestorm wave base.

6b Argillaceous wackestone topackstone

W-P peloids (3), skeletal debris (3),echinoderms (2), benthicforaminifera (2), Chofatella decipiens(2), porifera (2), oysters (2–3)

bioturbation Low energy deep subtidalplatform belowstorm wavebase.


FIG. 3.—Field pictures showing characteristics of the Jurf and the Qishn Formation dolomites. A) Wavy microbial mats with desiccation cracks (arrows). B) Thinsection of microbial mats under plane-polarized light showing wavy crinkly laminations (arrow) and folded thin layers of organic matter (om). The porosity is coloredwith blue dye. C) Qishn Formation dolomite showing calcite nodules (arrows) interpreted to be a replacement of anhydrite nodules. D) Extensive calcite cementation(arrow) and high degree of outcrop weathering in the Middle Jurf Formation.

FIG. 4.—XRD and petrographic results forthe Jurf Formation showing the stratigraphicdistribution of various mineral phases andpetrographically distinct dolomite phases.


Strontium isotope analyses (87Sr/86Sr) were carried out on 16 bulkdolomite samples, formed of more than 95% dolomite in the carbonatefraction and less than 30% of insoluble residue, and one bulk limestonesample. The insoluble residue is calculated by measuring the difference inmass of a sample before and after acidification by a 5% HNO3 solutionand heating at 80uC for one hour. Strontium was separated from thesolutions using Eichrom Sr-spec resin, and analyzed on a VG354 thermalionization mass spectrometer at Royal Holloway University of London.Samples were loaded on single Re filaments with a TaF emitter and run

using the multidynamic procedure of Thirlwall (1991). All values werenormalized to SRM 987 (0.710248). The standard error (2 S.E.) onindividual analysis was between 0.000010 and 0.000012.

Concentrations of Sr, Mn, and Fe in dolomite were determined in thinsections (164 acquisition points) by electron-probe microanalysis (EPMA)on a CAMECA SX-100 instrument equipped with five wavelength-dispersive X-ray spectrometers (WDS) at Service Microsonde Sud,Universite Montpellier II. The analyses were done with 20 kV acceler-ating voltage, a focused beam of 15 mm, and counting times of 20–30 s.

FIG. 5.—Main dolomite types (scale bar is 100 mm). A) Fine crystalline dolomite (D1) showing: 1) diagenetic quartz and 2) corroded quartz grains. B) D1 vuggydolomite showing molds of skeletal grains. C) D1 Qishn Formation under CL, note the homogeneity of luminescence. D) Stylolite transecting D1 dolomite phase.E) Medium crystalline dolomite D2 under CL, note the patchy luminescence. F) Coarse crystalline dolomite D3 on pore rims showing CL zonations, the void is filled withnonluminescent calcite cement with bright orange zonations. G) Isolated dolomite rhomb D4 along stylolite.


Concentrations were obtained from raw intensities using the ‘‘X-PHI’’quantification procedure (Merlet 1994). Natural minerals, syntheticoxides, and pure metals were used as standards. Average detection limitsat 95% confidence level are as follows: Fe 5 55 ppm; Mn 5 90 ppm;Sr 5 128 ppm. Due to the high detection limits of Sr and Mn in EPMA,elemental concentrations were also obtained on solutions (prepared frombulk samples) by inductively coupled plasma atomic emission spectro-scopy (ICP-AES) at the Natural History Museum London. Aliquots of100 mg of 12 bulk sample powders were dissolved in 20 mL of a 5%HNO3 solution and then heated to about 80uC for one hour. Values wereobtained for samples composed of more than 95% dolomite in thecarbonate fraction and less than 15% of insoluble residue because thepresence of Sr-rich impurities such as gypsum, anhydrite, and saline fluidinclusions complicates interpretations. Analytical precision is ca. 10% fordolomite and limestone standards.

RESULTS

Facies Types

Facies were defined based on the observation of sedimentary structuresand faunal assemblages at outcrop and in thin section (Table 1, Fig. 3). Adetailed description of every facies besides the siliciclastic–carbonatefacies and the microbial laminite facies can be found in Sena and John(2013). XRD analyses (Fig. 4) indicate that, apart from dolomite andcalcite, Jurf sedimentary rocks contain quartz, low amounts of clay andgypsum, and trace amounts of orthoclase, plagioclase, and pyrite.

Petrographic and Stratigraphic Character of the Dolomite

D1: Fine Crystalline Dolomite.—The Lower and Middle Jurf Formationlimestones and two beds in the Qishn Formation were replaced by non-mimetic fabric-destructive fine crystalline dolomite, hereafter named D1(Fig. 5A–D). D1 dolomites present planar-s to planar-e textures with crystalsize ranging between , 4 mm to 50 mm and usually exhibit a unimodalcrystal size distribution. D1 dolostones contain vuggy and moldic porosity(Fig. 5B) with variable amounts of intercrystalline matrix porosity(Fig. 5C). D1 exhibits dull orange to dull brown uniform luminescence(Fig. 5C). Stylolites and stylolitic seams crosscut D1 (Fig. 5D).

D2: Medium Crystalline Dolomite.—Three out of the 105 samples of theJurf Formation are characterized by non-mimetic fabric-destructivemedium crystalline dolomites (65 mm in average), hereafter named D2.D2 presents planar-e textures, and D2 dolostones have mostly vuggyporosity with little intercrystalline matrix porosity (Fig. 5E). D2dolomites are zoned in plane light and present patchy dull and brightorange luminescence. D2 dolomites occur locally and cannot becorrelated between the sections (Fig. 4).

D3: Coarse Crystalline Dolomite.—Dolomite crystals 60 to 200 mm insize, hereafter named D3, nucleate on the rims of pores filled with calcitecement and grow towards the center of the pore (Fig. 5F). D3 presentsalternated dull brown and dull orange crystal growth zones and dullorange and bright orange zonation.

D4: Isolated Dolomite Rhombs in Limestone.—Isolated euhedraldolomite rhombs (D4), commonly 100 to 250 mm across, occur floatingin the matrix of mudstone and wackestone and can be found alongstylolites (Fig. 5G). Isolated dolomite rhombs have cloudy centers andpatchy dull and bright luminescence.

Calcite Cement.—In 25 out of 105 samples of the Jurf Formation theintercrystalline, vuggy, and moldic porosity of D1 dolostones is filled withcalcite cement. Three types of calcite luminescence patterns wereidentified: patchy dull orange (Fig. 6A), nonluminescent with dull andbright orange zonation (Figs. 5F, 6B), and uniform nonluminescent.

Molds of isolated dolomite rhomb are found along stylolites and in themicritic matrix (Fig. 7A, B), and calcite cement with patchy dull orangeluminescence is found in calcitized D1 and D3 dolomite rhombs (Fig. 7D,E). Crosscutting relationships between different calcite cement phasescould not be established petrographically, given the lack of diagnosticstructures such as veins.

Stratigraphic Distribution of Mineralogies and Textures

The lower part of the Jurf Formation (lower 6 meters from the top ofthe Permian Gharif Formation) was pervasively replaced by D1dolomite and small amounts of calcite cement (1 to 15% of calcite persample). The middle part of the Jurf Formation mostly comprises

FIG. 6.—A, B) Main calcite cement types (scale bar is 100 mm). Fine crystalline dolomite D1 and coarse crystalline dolomite D3 (arrows) showing patchy dull and brightorange CL pattern. Note the bright rim around D3 dolomite (arrow in Part B). Calcite cement CL patterns show evidence for several phases of dissolution–precipitation.


calcitic dolostones composed of D1 and D3 dolomite rhombs, andcalcite forms 15% to 80% of the carbonate fraction. The upper part ofthe Jurf Formation comprises partially dolomitized limestones contain-ing less than 10% of D4 dolomite. The contact between the Jurfdolostone and the Jurf partially dolomitized limestone is sharp andcorresponds to a bedding plane (Fig. 8). The middle part of the JurfFormation forms extensively weathered gentle slopes at outcrop, and thetop-Jurf unconformity forms a table top at outcrop with evidence forextensive calcite cementation. Hence, two vertical trends independent of

facies occur upsection with increasing distance from the GharifFormation: D1 dolomite decreases, whereas calcite cement and D3and D4 dolomite increases (Figs. 4, 8).

Dolomitization in the Qishn Formation is limited to two beds 0.5 to 1 mthick, laterally continuous across 1 km to more than 2 km, and partialdolomitization occurs immediately on top of, and below, the dolostones(Fig. 8). The vertical contact between dolostone and dolomitic limestone issharp, and the lateral transition between the dolostone and the limestonespreads over several meters.

FIG. 7.—Main petrographic evidence for dolomite dissolution and dedolomitization (scale bar is 100 mm). A) Isolated dolomite rhomb molds (D4) along an openfracture. B) D4 dolomite rhomb molds: 1) in the micritic matrix and 2) along stylolites. C) Intercrystalline porosity of fine crystalline dolomite (D1) filled with calcitecement. D) Coarse crystalline dolomite (D3) rhomb partially replaced by calcite cement dyed in pink. E) Phantom of D3 dolomite rhomb replaced with calcite cementwith zoned CL pattern in vug filled with calcite.


FIG. 8.—Logged sections of the Jurf and Qishn formations showing lateral and vertical distribution of facies, dolostones, and dolomitic limestones. The satellite photo(Google Earth) of the outcrop on the upper left corner shows the location of the logged sections. Refer to Table 1 for the facies scheme.


GEOCHEMISTRY

Oxygen and Carbon Isotopic Ratios and Clumped Isotopes

The average d18O values for Qishn Formation and Jurf Formationdolostones are similar, and the d18O and d13C values for the dolostonesare enriched by 3.5% to 2.9% and 0.5% to 0%, respectively, whencompared to micrite in laterally equivalent formation limestones (Fig. 9,Table 2). The calcitic dolostones show a positive correlation betweenincreased calcite content and more depleted d18O and d13C values.

The clumped-isotopes D47 ARF values and calculated temperatures forthe Jurf and Qishn dolostones are within error of each other (i.e.,44 6 3uC, 1 S.D., Table 3). The values of d18Ofluid calculated usingthe temperature relationship of Vasconcelos et al. (2005) relatingtemperature, d18Ofluid and d18Odolomite range from 2.8% to 3.5 6 0.4%(Table 3).

Strontium Isotopes

87Sr/86Sr values for Qishn Formation dolostones (on average 0.70756,n 5 3) are lower than 87Sr/86Sr values for Jurf Formation dolostones (onaverage 0.70782, n 5 13). The 87Sr/86Sr values of Qishn Formationdolostones are similar to Barremian seawater (0.70740–0.70750), and87Sr/86Sr values of Jurf Formation dolostones are within the range ofPermian seawater values (0.7073–0.7083) (Veizer et al. 1999). In section G,the amount of radiogenic Sr decreases upsection (Fig. 10A). Weakcorrelations exist between the insoluble residue and the 87Sr/86Sr values forsections J (R2 5 0.50), G (R2 5 0.30), and I (R2 5 0.37), but there is anoverall increase in 87Sr/86Sr with increased content of insoluble residue forsections I and G (Fig. 10B).

Major and Trace Elements

The Qishn dolostones (n 5 3) have 54 to 56 mol percent CaCO3,indicating that the Qishn dolostones are slightly less stoichiometric thanthe Jurf dolostones with 50 to 54 mol % CaCO3 (Fig. 11A, Table 2). TheSr concentrations in the Lower Jurf Formation are characterized by alinear decrease upsection by a factor of 1.2 to 1.6 and a lateral decrease byca. 30 6 10 ppm from sections I and J to section G (Fig. 11B). QishnFormation dolostones have generally higher Sr concentrations comparedto the Jurf Formation dolostones, and the lateral variation of Sr along asingle bed is 70 ppm.

Jurf and Qishn dolostones have distinct Fe and Mn concentrations(Table 2). There is a nearly one order of magnitude difference in Fe

concentrations between the Qishn Formation dolostones (average of964 ppm) and the Jurf Formation dolostones (average of 6011 ppm).The same is true for Mn concentrations, with Qishn Formationdolostones having on average 48 ppm versus 271 ppm in the JurfFormation dolostones. Fe concentration is highest in sandy dolostonesand ranges between 2000 and 7000 ppm in Jurf Formation dolostones(Fig. 11C).

INTERPRETATION AND DISCUSSION

Paragenesis and Timing of Diagenetic Events

The Jurf and Qishn Formation D1 dolomites are characterized by smallcrystal sizes indicating multiple-site nucleation and relatively rapid crystalgrowth on mud-rich limestone precursors (Sibley 1982). D1 crystal sizedistribution does not follow any stratigraphic trend, and the homogeneousCL patterns suggest only one phase of dolomite precipitation. On thecontrary, D2 dolomite crystals present patchy CL patterns interpreted tobe indicative of multiple dolomite dissolution–precipitation phases.

Matrix dolomitization (D1 precipitation) and the development ofvuggy and moldic porosity must have preceded precipitation of D3,inasmuch as this latter dolomite phase occurs on the rims of vugs andmolds. It is difficult to determine whether D3 precipitated as dolomitecement or as a replacement of a precursor vug-filling calcite cement. TheCL zonations of the calcite are not disrupted by D3, and the CLzonations of D3 dolomite rhombs suggest that they nucleated on the porerim and grew towards the center of the vug. Therefore, we hypothesizethat D3 is a cement that precipitated in empty vugs before calcitecementation. The D4 dolomite phase occurs in the limestone matrix butalso in fractures and stylolites, suggesting that it precipitated aftercompaction in a burial setting.

Phantoms of dolomite rhombs found in calcite cement suggest thatcalcite replaced dolomite and that D1, D3, and D4 underwent a phase ofdedolomitization. Calcite cements with alternating dark, bright, andmoderately luminescent zones may result from precipitation from porewaters with fluctuating chemistry or in fluctuating redox conditions(Machel 2000), typical of the meteoric domain. The depleted carbon andoxygen isotopic ratios of the dedolomites also suggest that the diageneticfluid involved in the dedolomitization was of meteoric origin (Lohmann1988). Furthermore, high amounts of calcite cementation associated withhigh degrees of outcrop weathering, as well as dolomite dissolution alongstylolites and fractures, both suggest that most of the dolomite dissolutionand calcite precipitation were late processes resulting from circulation of

FIG. 9.—Cross-plot of dolomite d13C andd18O values for Jurf and Qishn Formationdolomite. The range of d13C and d18O values forthe Qishn and Jurf Formation limestone (Senaand John 2013) and for Cretaceous marineseawater (Veizer et al. 1999) is indicated.


meteoric fluids. This intense phase of meteoric diagenesis could havetaken place during the wet Pleistocene period, when the formation wascropping out, similar to what has been suggested for the OmanMountains by Vandeginste and John (2012).

Origin and Circulation Mechanisms of the Dolomitizing Fluids

Results show that D1 dolomite in both the Jurf and Qishn formationswas non-mimetic fabric destructive and replaced several strata ofperitidal, subtidal, and sandy limestones. The oxygen isotopic ratios ofthe fluids that precipitated the D1 phase range between +2.8% to +3.5%SMOW. These values are higher than the reported Cretaceous seawatervalues (22.6% to 1.2% SMOW, Veizer et al. 1999), whereas the d13Cvalues of D1 fall within the range of expected values for Cretaceousseawater (Veizer et al. 1999). This suggests that the dolomitizing fluid forthe Jurf and Qishn Formation was Cretaceous seawater that underwentsome degree of evaporation.

Independent arguments in support of the interpretation of evaporativeconditions are: 1) the microbial laminite facies at the base of the JurfFormation and the presence of abundant miliolids in beds above andbelow Qishn dolostones is evidence of restricted conditions with highsalinities (Immenhauser et al. 2004); 2) calcite nodules present in theQishn dolostone have a size (2 to 10 cm long) and regular shape (sphericalto elliptical) suggestive of anhydrite relicts. The deposition of restrictedsedimentary facies and the presence of anhydrite relicts indicate that thedolomitizing fluid could occasionally reach hypersalinity. Becauseevaporites have not been observed in the studied supratidal paleoenvir-onments, it is suggested that seawater salinity was never sufficient todevelop an evaporative lagoon.

Processes for the near-surface circulation of slightly evaporated seawaterdriven by brine density and/or sea-level fluctuations include reflux (Adamsand Rhodes 1960), evaporative pumping (Hsu and Siegenthler 1969), andmarine recharge (Carballo et al. 1987). Reflux has proven to occur whereseawater is below gypsum saturation (Simms 1984; Melim and Scholle2002) and therefore possible to occur in carbonate sequences lackingevaporites such as the Jurf and Qishn formations (Fig. 12A).

The link between dolomite occurrence and the presence of microbialmats suggests that dolomite precipitation might have been induced bybacterially mediated reduction of seawater sulfate (Vasconcelos andMcKenzie 1997; van Lith et al. 2003; Bontognali et al. 2010; Krause et al.2012) and that dolomite may have precipitated from normal seawater.This model does not require huge volumes of fluids to be pumped throughthe carbonate platforms, as would be the case for dolomitization byreflux, and would explain the limited distribution and strataboundgeometry of the Qishn Formation dolomite.

Mechanism and Timing of Burial Diagenesis

The temperature of Jurf and Qishn D1 dolomite formation (44 6 3uC)indicate that the early dolomites re-equilibrated to fully ordereddolomites during shallow burial, as shown in Vasconcelos et al. (2006).The relatively high Fe, Mn concentrations and high 87Sr/86Sr ratios in theJurf Formation dolostones suggest that the burial fluid interacted withnoncarbonate phases (Banner et al. 1988). The fact that the highest Feconcentrations are found in the Jurf Formation sandy dolostones suggeststhat the local siliciclastic grains such as Rb- and Fe-rich clay minerals,feldspars, and micas sourced from the Permian clastics were a source ofFe during dolomitization. The Fe trapped in microbial mats is anotherpossible local source of this element (Krumbein et al. 2003). The decreaseupsection of 87Sr/86Sr, Sr, and Fe in the Jurf Formation suggests that theburial fluids were expelled upwards from the underlying Permian sandsduring sediment compaction. By contrast, the 87Sr/86Sr values of theQishn Formation dolostones are within the range of 87Sr/86Sr values ofT

AB

LE

2.—

Pet

rogr

aphi

can

dge

oche

mic

alch

arac

teri

stic

sof

fine

crys

tall

ine

dolo

mit

e(D

1)an

dco

arse

crys

tall

ine

dolo

mit

e(D

3).D

olom

ite

tex

ture

s,C

Lpa

tter

nsan

dcr

ysta

llen

gth

are

indi

cate

d.N

ote

that

Mn

cont

ent

for

the

Qis

hnF

orm

atio

nan

dS

rco

nten

tw

ere

mea

sure

dw

ith

ICP

-AE

Sw

here

asal

lot

her

elem

enta

lda

taw

asm

easu

red

wit

hE

PM

A.

Mea

sure

men

tsm

ark

edw

ith

a*

are

from

bulk

sam

ples

;n

isth

enu

mbe

rof

sam

ples

;S

Dis

one

stan

dard

devi

atio

n.

Fo

rmat

ion

Do

lom

ite

Ph

ase

Do

lom

ite

Tex

ture

CL

Pat

tern

Cry

stal

Len

gth

(mm

)d1

3C

do

l*%

(n,

S.D

.)d1

8O

do

l*%

(n,

S.D

.)87S

r/86S

r*(n

)F

ep

pm

(n)

Sr

pp

m(n

)M

np

pm

(n)

CaC

O3*

mo

l%

Qis

hnF

mD

1eu

hed

ral

du

lld

ark

bro

wn

bro

wn

30–6

02.

5(3

,0.

2)2

1.2

(3,

0.2)

0.70

755–

0.70

758

(3)

611–

1317

(2)

105–

173

(2)

31–6

9(2

)53

.7–5

6.1

Upp

erJ

urf

Fm

D3

euh

edra

lin

styl

oli

ted

ull

bri

ght

pat

chy

ora

nge

100–

200

--

-84

0(1

),

det

lim

83(1

)-

Mid

dle

Jur

fF

mD

3eu

hed

ral

invu

gri

mzo

ned

du

llo

ran

gean

db

row

n60

–200

--

-19

28–3

323

(3)

,d

etli

m18

6–24

4(2

)-

D1

sub

hed

ral

and

euh

edra

ld

ull

tob

righ

to

ran

gean

dzo

ned

on

vug

rim

s

,4

to30

22.

4(1

6,0.

9)2

3.7

(16,

1.7)

-47

86–6

609

(3)

,d

etli

m19

2–41

9(3

)50

.9–5

4

Low

erJ

urf

Fm

D1

sub

hed

ral

and

euh

edra

ld

ull

ora

nge

,4

to30

20.

4(8

1,0.

6)2

1.2

(81,

0.7)

0.70

762–

0.70

810

(13)

2376

–950

7(8

)85

–135

(8)

130–

377

(8)

50.3

–51.

8


Cretaceous seawater (Veizer et al. 1999), and the low Mn and Feconcentrations suggest that the burial fluid did not interact withnoncarbonate phases. This suggests that burial fluids also circulatedlaterally into less rapidly compacting sediments (Fig. 12B).

Because compaction expels a single pore volume of fluid from thesource rocks, it cannot generate sufficient fluid flow to explain the originof massive dolomite bodies (Demming et al. 1990; Machel and Calvell1999). It can, however, explain the precipitation of D3 and D4 dolomitephases occurring in vugs, fractures, stylolites, and limestone matrix ofsedimentary rocks adjacent to the dolostones (Fig. 12B).

Jurf and Qishn Dolomite in the Context of Regional and GlobalDolomitization Models

The study of synchronism of dolomitizing events in the Phanerozoic(Given and Wilkinson 1987; Sun 1994) and the Cenozoic (Budd 1997) ledto the formulation of two hypotheses for the forcing mechanism ofdolomitization, one related to glacio-eustatic change and the other toglobal climatic fluctuations.

The first hypothesis stipulates that the most effective dolomitizingenvironments are those that circulate large volumes of fluids, a conditionthat is most likely to be met in hydrological systems located near meansea level during prolonged periods of time (Sibley 1991). This implieslarge volumes of dolomite being formed either during long highstands orlong lowstands. Furthermore, a genetic relation between dolomite andsea-level history has been identified based on the relation betweendolomite abundance and maximum continental flooding (Given andWilkinson 1987). It is thought that seawater dolomitization may beenhanced during times of global transgression characterized by higheratmospheric pCO2, lower oceanic CO3

22 concentrations, and lowercalcite saturation state (Machel and Mountjoy 1986). The secondhypothesis considers dolomitization to be associated with global climaticaridity (Sibley 1980; McKenzie 1991) and considers salinity as a criticalkinetic factor (Sun 1994) based on the observation that most large-scaledolomitization had an origin related to evaporated seawater.

The dolomite distribution in the Lower Cretaceous of Oman supportsboth the first hypothesis (most of the dolomitization occurs during thefirst transgressive phase of a low-frequency cycle and is associated withcontinental flooding, low accommodation, and aggradation of peritidalmicrobial mats) and the second hypothesis (dolomitization occurs in aperitidal environment characterized occasionally by hypersalinities).Nevertheless, although both the Jurf Formation and the QishnFormation are composed mainly of peritidal facies deposited underslightly elevated salinities (Sena and John 2013), the volume of dolomitegenerated by similar near-surface dolomitization models differs. Asignificant volume of dolomite is generated in the Jurf Formation, butdolomitization is restricted to only two continuous beds in the QishnFormation. Because of the absence of positive correlation between watersalinities and dolomitization potential, we hypothesize that salinity of thefluids did not exert a primary control on dolomitization.

One possible explanation for the disparity in dolomite volumes betweenthe Jurf and Qishn formations is to hypothesize that the Permian sandsbelow the Jurf Formation had an influence on the amount of dolomite

TABLE 3.—Clumped isotopes results, calculated temperatures of dolomite formation and calculated d18Ofluid of the dolomitizing fluid for the Jurf and Qishnfine crystalline dolomite D1. ARF: Absolute reference frame.

Formation Sample Name D47 ARF 6 1 S.E. Temperature in uC (Ghosh et al., 2006) d18O (VPDB) d18Ofluid (SMOW) (Vasconcelos et al., 2005)

Qishn Formation E36 0.623 6 0.004% 45 6 2uC 20.72% 3.3 6 0.4%Jurf Formation G5 0.621 6 0.004% 46 6 2uC 21.46% 2.8 6 0.4%

I4 0.628 6 0.005% 44 6 2uC 20.21% 3.5 6 0.4%Average temperature for all samples: 44 ± 3uC

FIG. 10.—A) 87Sr/86Sr for D1 dolomites plotted against stratigraphic distancefrom the Gharif Formation. B) Cross-plot of the insoluble residue and 87Sr/86Sr forD1 dolomites in the Jurf Formation.


generated by acting as a permeable bed and by playing a major role onmaintaining active fluid circulation during early dolomite formation. Asecond explanation could be related to the different sea-level histories ofthe Jurf and Qishn formations, as suggested by the numerousdiscontinuity surfaces that characterize the peritidal environment of theQishn Formation (Sattler et al. 2005; Sena and John 2013). Short-termsea-level falls in the Qishn Formation could have been more prolongedand resulted in the flushing of the platform top by meteoric waters,decreasing the Mg/Ca ratio of seawater to levels insufficient fordolomitization. In addition, higher subsidence rates in the outer platformresulted in supratidal environments existing for a shorter portion of cycleduration than in the inner platform, thus limiting dolomitization.

The main difference between the depositional environment of the LowerJurf and Qishn formations is the abundance of stromatolitic and microbialmat-like lamination in the former. The abundance of microbial mats in theJurf Formation may have markedly increased its dolomitization potential,as observed in the Abu Dhabi sabkha (Bontognali et al. 2010). It alsowould explain why dolomite is found in the transgressive systems tract ofthe Jurf Formation when the conditions for the near-surface dolomitiza-tion models mentioned above are not optimal. The decrease in abundanceof microbial mats (and hence dolomite) is suggested to be related to achange from an arid to a humid climate during the Barremian–Aptian, asdescribed by Puceat et al. (2003) and Steuber et al. (2005). The presumablymore humid climate would have prevented the development of bothmicrobial mats and evaporative conditions during the highstands.

Due to the lack of regional data, the areal extent of the Jurf Formationdolomite is difficult to ascertain. Jurf Formation dolostones occur in WadiJarrah, 60 km north of Wadi Baw (Immenhauser et al. 2004), suggesting that

dolomitization was a regional process. The supratidal zone that developedduring tidal-flat progradation on the Lower Cretaceous epeiric platformwould likely have been orders of magnitude more laterally extensive than inmodern sabkha environments (80 to 130 km maximum length (Purser 1973))given the low gradient of epeiric platforms (, 0.01 m/km) compared tomodern slopes (0.3 to 0.4 m/km) in Persian Gulf (Patterson and Kinsman1981). From the study of the controls on dolomite distribution, we suggestthat dolomitization was probably not effective in the Oman Interior basinsfor time-equivalent units because the presence of a long-lasting peritidalenvironment (found only around the Haushi-Huqf paleohigh) is proposedhere to be essential for the development of microbial mats and evaporativeconditions, with both controlling the dolomitization potential.

CONCLUSION

Peritidal and subtidal facies with fine crystalline dolomite in the JurfFormation occur in a body ten meters thick and at least tens of kilometerswide dolomite. In the Qishn Formation, however, dolomite replacementis restricted to two tidal-flat beds continuous over at least 1 km to morethan 2 km.

Petrographic relations, stratabound geometry, and distribution ofdolomite mainly in facies rich in microbial mats suggest that dolomiti-zation was an early, near-surface process. The dolomitizing fluid wasslightly evaporated Cretaceous seawater that circulated through thecarbonate platform through density-driven and sea-level-fluctuation-driven flow such as reflux, evaporative pumping, and marine recharge.The close association of dolomite and microbial-mat facies indicate thatdolomite precipitation from normal seawater might have also been

FIG. 11.—A) CaCO3 mol percent in finecrystalline dolomite (D1) determined with XRDdata. B) Sr content in D1 measured with EPMAand ICP-AES. C) Fe content in D1 and inisolated dolomite rhombs (D4) measured withEPMA. Standard deviations of Fe elementalconcentrations per thin section are shown. Alldata are plotted against the distance from the topof the Gharif Formation. Uncertainty of 2.5 me-ters in the correlation to section G is indicated.


possible through the lowering of kinetic barriers to dolomite precipitationby bacteria. The paleo-temperature suggested by clumped isotopes (44uC)and elevated Sr, Fe, Mn, and 87Sr/86Sr are all indicative that the earlydolomite re-equilibrated with shallow-burial fluids. The presence of smallvolumes of coarse crystalline dolomite in stylolites and fractures suggeststhat burial fluids circulated as a consequence of sediment compaction.

Comparison of the dolomite in the peritidal environment of the Jurfand Qishn formations suggests that 1) peritidal carbonates on an epeiricplatform can be dolomitized even if no abundant evaporitic facies arepresent, and 2) the development of microbial mats associated with climatearidity exerts a major control on early dolomite distribution.

ACKNOWLEDGMENTS

We gratefully acknowledge funding from the Qatar Carbonates andCarbon Storage Research Center (QCCSRC), provided jointly by QatarPetroleum, Shell, and the Qatar Science & Technology Park. The reviewers,Adrian Immenhauser and Bruce Wilkinson, are deeply thanked for theirinsightful comments. Fruitful discussions at an early stage of this work withConxita Taberner, Paul Wagner, Rob Forkner, Chris Nicholls, AndreaKnoerich, and Christian Tueckmantel were greatly appreciated. Simon Davis,Catherine Unsworth, Martin Gill, and Bernard Boyer are thanked for theirtechnical assistance in the laboratories. Anna Joy Drury, Simon Davis, ReniJoseph, and Francisco Pereira from Shuram Oil & Gas, L.L.C. are deeplythanked for their assistance in the field.

REFERENCES

ADAMS, J.E., AND RHODES, M.L., 1960, Dolomitization by seepage refluxion: AmericanAssociciation of Petroleum Geologists, Bulletin, v. 44, p. 1912–1920.

ALLISON, P.A., AND WELLS, M.R., 2006, Circulation in large ancient epicontinental seas:What was different and why?: Palaios, v. 21, p. 513–515.

BANNER, J.L., HANSON, G.N., AND MEYERS, W.J., 1988, Water–rock interaction historyof regionally extensive dolomites of the Burlington–Keokuk Formation (Mississip-pian): isotopic evidence, in Shukla, V., and Baker, P.A., eds., Sedimentology andGeochemistry of Dolostones: SEPM, Special Publication 43, p. 97–113.

BONTOGNALI, T.R.R., VASCONCELOS, C., WARTHMANN, R.J., BERNASCONI, S.M., DUPRAZ, C.,STROHMENGER, C.J., AND MCKENZIE, J.A., 2010, Dolomite formation within microbialmats in the coastal sabkha of Abu Dhabi (United Arab Emirates): Sedimentology, v. 57,p. 824–844.

BRISTOW, T.F., BONIFACIE, M., DERKOWSKI, A., EILER, J.M., AND GROTZINGER, J.P., 2011,A hydrothermal origin for isotopically anomalous cap dolostone cements from southChina: Nature, v. 474, p. 68–92.

BUDD, D.A., 1997, Cenozoic dolomites of carbonate islands: their attributes and origin:Earth-Science Reviews, v. 42, p. 1–47.

CARBALLO, J., LAND, L.S., AND MISER, D.E., 1987, Holocene dolomitization ofsupratidal sediments by active tidal pumping, Sugarloaf Key, Florida: Journal ofSedimentary Petrology, v. 57, p. 153–165.

CHOQUETTE, P.W., AND PRAY, L.C., 1970, Geologic nomenclature and classification ofporosity in sedimentary carbonates: American Association of Petroleum Geologists,Bulletin, v. 54, p. 207–244.

DEMMING, D., NUNN, J.A., AND EVANS, D.G., 1990, Thermal effects of compaction-driven groundwater flow from overthrust belts: Journal of Geophysical Research,v. 95, p. 6669–6683.

DENNIS, K.J., AND SCHRAG, D.P., 2010, Clumped isotope thermometry of carbonatites as anindicator of diagenetic alteration: Geochimica et Cosmochimica Acta, v. 74, p. 4110–4122.

FIG. 12.—A) Conceptual dolomitizationmodel for the formation of the Jurf Formationand Qishn Formation fine crystalline dolomiteD1. The schematic location of the sections andtheir spacing is shown for reference. B) Con-ceptual model for the re-equilibration of D1 andprecipitation of D3 and D4 dolomite phases.


DENNIS, K.J., AFFEK, H.P., PASSEY, B.H., SCHRAG, D.P., AND EILER, J.M., 2011, Definingan absolute reference frame for ‘‘clumped’’ isotope studies of CO2: Geochimica etCosmochimica Acta, v. 75, p. 7117–7131.

DICKSON, J.A.D., 1966, Carbonate identification and genesis as revealed by staining:Journal of Sedimentary Petrology, v. 36, p. 491–505.

DRAVIS, J.J., 1991, Carbonate petrography: update on new techniques and applications:Journal of Sedimentary Petrology, v. 61, p. 626–628.

DROSTE, H., 2010, Sequence-stratigraphic framework of the Aptian Shu’aiba Formationin the Sultanate of Oman, in Van Buchem, F.S.P., Al-Husseini, M.I., Maurer, F., andDroste, H., eds., Barremian–Aptian Stratigraphy and Petroleum Habitat of theEastern Arabian Plate: GeoArabia, Special Publication 4, p. 229–283.

DUNHAM, R.J., 1962, Classification of carbonate rocks according to their depositionaltexture, in Ham, W.E., ed., Classification of Carbonate Rocks: American Associationof Petroleum Geologists, Memoir, v.1, p. 108–121.

FILBRANDT, J.B., AL-DHAHAB, S., AL-HABSY, A., HARRIS, K., KEATING, J., AL-MAHRUQI,S., ISMAIL OZKAYA, S., RICHARD, P.D., AND ROBERTSON, T., 2006, Kinematicinterpretation and structural evolution of North Oman, Block 6, since the LateCretaceous and implications for timing of hydrocarbon migration into Cretaceousreservoirs: Geoarabia, v. 11, p. 97–140.

GHOSH, P., ADKINS, J., AFFEK, H., BALTA, B., GUO, W., SCHAUBLE, E., SCHRAG, D., AND

EILER, J., 2006, 13C–18O bonds in carbonate minerals: a new kind of paleotherm-ometer: Geochimica et Cosmochimica Acta, v. 70, p. 1439–1456.

GIVEN, R.K., AND WILKINSON, B.H., 1987, Dolomite abundance and stratigraphic age-constrains on rates and mechanisms of Phanerozoic dolostone formation: Journal ofSedimentary Petrology, v. 57, p. 1068–1078.

GRADSTEIN, F.M., OGG, J.G., AND SMITH, A.G., 2004, A Geological Timescale 2004:Cambridge U.K., Cambridge University Press, 589 p.

HARDIE, L.A., 1987, Dolomitization: a critical view of some current views: Journal ofSedimentary Petrology, v. 57, p. 166–183.

HSU, K.J., AND SIEGENTHLER, C., 1969, Preliminary experiments on hydrodynamicmovement of induced evaporation and their bearing on the dolomite problem:Sedimentology, v. 12, p. 11–25.

HUNTINGTON, K.W., EILER, J.M., AFFEK, H.P., GUO, W., BONIFACIE, M., YEUNG, L.Y.,THIAGARAJAN, N., PASSEY, B., TRIPATI, A., DAERON, M., AND CAME, R., 2009, Methodsand limitations of ‘‘clumped’’ CO2 isotope (D 47) analysis by gas-source isotope ratiomass spectrometry: Journal of Mass Spectrometry, v. 44, p. 1318–1329.

IMMENHAUSER, A., HILLGARTNER, H., SATTLER, U., BERTOTTI, G., VAN DER KOOIJ, B.,VAN BENTUM, E., VAN KOPPEN, J., VERWER, K., IMMENHAUSER-POTTHAST, I.,SCHOEPFER, P., VAHRENKAMP, V., HOOGERDUIJN-STRATING, E., PETERS, J., HOMEWOOD,P., DROSTE, H.J., SWINKELS, W., STEUBER, T., MASSE, J.P., AND AL MASKERY, S.A.J.,2004, The Barremian–Lower Aptian Qishn Formation (Huqf Area, Oman): anoutcrop analogue for Kharaib/Shuaiba Subsurface Reservoirs: Geoarabia, v. 9,p. 153–194.

KRAUSE, S., LIEBETRAU, V., GORB, S., SANCHEZ-ROMAN, M., MCKENZIE, J.A., AND

TREUDE, T., 2012, Microbial nucleation of Mg-rich dolomite in exopolymericsubstances under anoxic modern seawater salinity: new insights into an old enigma:Geology, v. 40, p. 187–194.

KRUMBEIN, W.E., BREHM, U., GERDES, G., GORBUSHINA, A.A., LEVIT, G., AND PALINSKA,K.A., 2003, Biofilm, biodictyon, biomat microbialites, oolites, stromatolites,geophysiology, global mechanism, parahistology, in Krumbein, W.E., Paterson,D.M., and Zavarzin, G.A., eds., Fossil and Recent Biofilms: A Natural History ofLife on Earth: Dordrecht, Kluwer, p. 1–28.

LAND, L.S., 1980, The isotopic and trace element geochemistry of dolomite: the state ofart, in Zenger, D.H., Dunham, J.B., and Ethington, R.L., eds., Concepts and Modelsof Dolomitization: SEPM, Special Publication 28, p. 87–110.

LAND, L.S., 1985, The origin of massive dolomite: Journal of Geo-Science Education,v. 33, p. 112–125.

LAND, L.S., 1998, Failure to precipitate dolomite at 25 uC from dilute solution despite1000-fold oversaturation after 32 years: Aquatic Geochemistry, v. 4, p. 361–368.

LE BEC, A., 2003, Distribution et dynamique des systemes carbonates de la plate-formecretace inferieur du Sultanat d’Oman [unpublished Doctoral thesis]: Universite Michelde Montaigne (Bordeaux 3), Bordeaux.

LIPPMAN, F., 1973, Sedimentary carbonate minerals: New York, Springer-Verlag, 228 p.LOHMANN, K.C., 1988, Geochemical Patterns of Meteoric Diagenetic Systems and Their

Application to Studies of Paleokarst, in James, N.P., and Choquette, P.W., eds.,Paleokarst: New York, Springer-Verlag, p. 58–80.

LOOSVELD, R.J.H., BELL, A., AND TERKEN, J.J.M., 1996, The tectonic evolution ofinterior Oman: GeoArabia, v. 1, p. 28–51.

LUMSDEN, D.N., 1979, Discrepancy between thin-section and X-ray estimates ofdolomite in limestone: Journal of Sedimentary Petrology, v. 49, p. 429–435.

MACHEL, H.G., 2000, Dolomite formation in Caribbean islands: driven by platetectonics?!: Journal of Sedimentary Research, v. 70, p. 977–984.

MACHEL, H.G., AND CALVELL, R., 1999, Low-flux, tectonically-induced squeegee fluidflow (‘‘hot flash’’) into the Rocky Mountain Foreland Basin: Bulletin of CanadianPetroleurn Geology, v. 47, p. 510–533.

MACHEL, H.G., AND MOUNTJOY, E.W., 1986, Chemistry and environments ofdolomitization: a reappraisal: Earth-Science Reviews, v. 23, p. 175–222.

MCKENZIE, J.A., 1991, The dolomite problem: an outstanding controversy, in Muller,D.W., McKenzie, J.A., and Weissert, H., eds., Controversies in Modern Geology:London, Academic Press, p. 37–54.

MCKENZIE, J.A., HSU, K.J., AND SCHNEIDER, J.F., 1980, Movement of subsurface watersunder the sabkha, Abu Dhabi, UAE and its relation to evaporative dolomite genesis,in Zenger, D.H., Dunham, J.B., and Ethington, R.L., eds., Concepts and Models ofDolomitization: SEPM, Special Publication 28, p. 11–30.

MELIM, L.A., AND SCHOLLE, P.A., 2002, Dolomitization of the Capitan Formationforereef facies (Permian, west Texas and New Mexico): seepage reflux revisited:Sedimentology, v. 49, p. 1207–1227.

MERLET, C., 1994, An accurate computer correction program for quantitative electronprobe microanalysis: Mikrochimica Acta, v. 114–115, p. 363–376.

PATTERSON, R.J., AND KINSMAN, D.J.J., 1981, Hydrologic framework of a sabkha along theArabian Gulf: American Association of Petroleum Geologists, Bulletin, v. 66, p. 28–43.

PLATEL, J.P., DUBREUILH, J., LE METOUR, J., ROGER, J., WYNS, R., BECHENNEC, F., AND

BERTHIAUX, A., 1992, Explanatory Notes and Geological Map of Duqm and Madraca,sheet NE 40-03/07.

PUCEAT, E., LECUYER, C., SHEPPARD, S.M.F., DROMART, G., REBOULET, S., AND

GRANDJEAN, P., 2003, Thermal evolution of Cretaceous Tethyan marine watersinferred from oxygen isotope composition of fish tooth enamels: Paleoceanography,v. 18, p. 1029.

PURSER, B.H., 1973, The Persian Gulf: Holocene Carbonate Sedimentation andDiagenesis in a Shallow Epicontinental Sea: New York, Springer, 471 p.

RIES, A.C., AND SHACKLETON, R.M., 1990, Structures in the Huqf-Haushi Uplift, eastCentral Oman: Geological Society of London, Special Publication 49, p. 653–663.

ROSENBAUM, J., AND SHEPPARD, S.M., 1986, An isotopic study of siderites, dolomitesand ankerites at high temperatures: Geochimica et Cosmochimica Acta, v. 54, p. 603–610.

SATTLER, U., IMMENHAUSER, A., HILLGARTNER, H., AND ESTEBAN, M., 2005, Character-ization, lateral variability, and lateral extent of discontinuity surfaces on a carbonateplatform (Barremian to Lower Aptian, Oman): Sedimentology, v. 52, p. 339–361.

SENA, C.M., AND JOHN, C.M., 2013, Impact of dynamic sedimentation on faciesheterogeneities in Lower Cretaceous peritidal deposits of central east Oman:Sedimentology, v. 60, p. 1156–1183.

SIBLEY, D.F., 1980, Climatic control of dolomitization, Seroe Domi Formation(Pliocene), Bonaire, N.A., in Zenger, D.H., Dunham, J.B., and Ethingtun, R.L.,eds., Concepts and Models of Dolomitization, SEPM, Special Publication 28, p. 247–258.

SIBLEY, D.F., 1982, The origin of common dolomite fabrics: Clues from the Pliocene:Journal of Sedimentary Petrology, v. 52, p. 1087–1100.

SIBLEY, D.F., 1991, Secular changes in the amount and texture of dolomite: Geology,v. 19, p. 151–154.

SIBLEY, D.F., AND GREGG, J.M., 1987, Classification of dolomite rock textures: Journalof Sedimentary Petrology, v. 57, p. 967–975.

SIMMS, M., 1984, Dolomitization by groundwater flow systems in carbonate platforms:Gulf Coast Association of Geological Societies, Transactions, v. 34, p. 411–420.

STEUBER, T., RAUCH, M., MASSE, J.-P., GRAAF, J., AND MALKOC, M., 2005, Low-latitudeseasonality of Cretaceous temperatures in warm and cold episodes: Nature, v. 437,p. 1341–1344.

SUN, S.Q., 1994, A reappraisal of dolomite abundance and occurrence in thePhanerozoic: Journal of Sedimentary Research, v. 64, p. 396–404.

THIRLWALL, M.F., 1991, Long-term reproducibility of multicollector Sr and Nd isotoperatio analysis: Chemical Geology: Isotope Geoscience Section, v. 94, p. 85–104.

VAHRENKAMP, V.C., AND SWART, P.K., 1990, New distribution coefficient for theincorporation of strontium into dolomite and its implications for the formation ofancient dolomites: Geology, v. 18, p. 387–391.

VAN BUCHEM, F., PITTET, B., HILLGARTNER, H., GROETSCH, J., AL MANSOURI, A.I.,BILLING, I.M., DROSTE, H.J., AND OTERDOOM, W.H., 2002, High-resolution sequencestratigraphic architecture of Barremian–Aptian carbonate systems in northern Omanand the United Arab Emirates (Kharaib an Shu’aiba Formations): Geoarabia, v. 7,p. 461–500.

VANDEGINSTE, V., AND JOHN, C.M., 2012, Influence of climate and dolomite compositionon dedolomitization: insights from a multi-proxy study in the central OmanMountains: Journal of Sedimentary Research, v. 82, p. 177–195.

VAN LITH, Y., WARTHMANN, R., VASCONCELOS, C., AND MCKENZIE, J.A., 2003, Sulphate-reducing bacteria induce low-temperature Ca-dolomite and high Mg-calcite forma-tion: Geobiology, v. 1, p. 71–79.

VASCONCELOS, C., AND MCKENZIE, J.A., 1997, Microbial mediation of modern dolomiteprecipitation and diagenesis under anoxic conditions (Lagoa Vermelha, Rio deJaneiro, Brazil): Journal of Sedimentary Research, v. 67, p. 378–390.

VASCONCELOS, C., MCKENZIE, J.A., WARTHMANN, R., AND BERNASCONI, S.M., 2005,Calibration of the d18O paleothermometer for dolomite precipitated in microbialcultures and natural environments: Geology, v. 33, p. 317–320.

VASCONCELOS, C., WARTHMANN, R., MCKENZIE, J.A., VISSCHER, P.T., BITTERMANN, A.G.,AND VAN LITH, Y., 2006, Lithifying microbial mats in Lagoa Vermelha, Brazil: ModernPrecambrian relics?: Sedimentary Geology, v. 185, p. 175–183.

VEIZER, J., ALA, D., AZMY, K., BRUCKSCHEN, P., BUHL, D., BRUHN, F., CARDEN, G.A.F.,DIENER, A., EBNETH, S., GODDERIS, Y., JASPER, T., KORTE, C., PAWELLEK, F., PODLAHA,

O.G., AND STRAUSS, H., 1999, 87Sr/86Sr, d13C and d18O evolution of Phanerozoic

seawater: Chemical Geology, v. 161, p. 59–88.VISSER, W., 1991, Burial and thermal history of Proterozoic source rocks in Oman:

Precambrian Research, v. 54, p. 15–36.

Received 20 January 2013; accepted 31 March 2014.


dolomitization of lower cretaceous peritidal … · 2015. 12. 14. · using the sibley and gregg...

Documents