Journal of Sedimentary Research, 2018, v. 88, 679–695

Research Article

DOI: http://dx.doi.org/10.2110/jsr.2018.35

STUDY OF AN ORDOVICIAN CARBONATE WITH ALTERNATING DOLOMITE–CALCITE LAMINATIONS

AND ITS IMPLICATION FOR CATALYTIC EFFECTS OF MICROBES ON THE FORMATION OF

SEDIMENTARY DOLOMITE

YIHANG FANG AND HUIFANG XU

NASA Astrobiology Institute, Department of Geoscience, University of Wisconsin–Madison, Madison, Wisconsin 53706, U.S.A.

e-mail: hfxu@geology.wisc.edu

ABSTRACT: The mechanism of sedimentary dolomite formation has puzzled the geology community for more than acentury. Within the past several years, successful synthesis of disordered dolomite under ambient conditions usingabiotic materials derived from microbial organisms such as polysaccharides and exopolymeric substances (EPS) hasbeen reported. The success in laboratory experiments has driven this study to find evidence in natural ancientcarbonate samples that correlate dolomite formation and the presence of organic matter. A micro-laminatedcarbonate with alternating dolomite–calcite layers from the mid-Lower Ordovician St. Paul Group from the CentralAppalachians in southern Pennsylvania was examined using optical microscopes, X-ray diffraction (XRD), scanningelectron microscopy (SEM) with X-ray energy-dispersive spectroscopy, electron microprobe analysis (EPMA),scanning transmission electron microscopy (STEM), laser-induced fluorescence (LIF) imaging, short-wave infrared(SWIR) imaging, and X-ray fluorescence (XRF) imaging. The sample is composed mainly of two types of layers.Dolomite-dominated layers are darker in color, generally thinner, and contain detrital minerals such as quartz andfeldspar. In contrast, calcite-dominated layers are lighter in color, thicker, and contain less detrital mineralssupported by microcrystalline calcite matrix. In situ XRD, LIF, XRF, and SWIR results show that organic remnantsare enriched in the dolomite layers. The coincided spatial distribution confirmed a positive correlation betweendolomite and organic matter, and hence provide evidence for microbial-EPS-catalyzed formation of sedimentarydolomite.

INTRODUCTION

Dolomite is one of the most common minerals in sedimentary rocks,

ranging from Archean to Holocene. Although dolomite is common in the

ancient rock record, formation of dolomite is scarce in modern sedimentary

settings (Hardie 1987; Warren 2000). Despite its widespread distribution

and economic value as water aquifers and petroleum reservoirs (Bereskin

et al. 2004; Braithwaite et al. 2004; Sonnenberg et al. 2011;), the formation

mechanism of sedimentary dolomite is still not well understood and is

known as the ‘‘dolomite problem’’ (Van Tuyl 1914; Zengler et al. 1980;

Tucker and Wright 1990; Land 1998; Kaczmarek et al. 2017; Petrash et al.

2017). Sedimentary dolomite is assumed to form in an environment with a

high Mg:Ca ratio such as seawater, some lacustrine water, and basinal

brines (Machel 2004). Synthesizing experiments with oversaturated

magnesium solution and durations of 32 years have shown that formation

of dolomite is not solely controlled by Mg/Ca ratio and salinity (Land

1998). The alternating Ca and Mg layers along the c axis in the dolomite

structure reduces the symmetry from R3c as in calcite and magnesite to R3.

However, many dolomites formed in modern settings and those

synthesized in the laboratory under ambient conditions are weakly ordered

or disordered, as evidenced by ‘‘b’’ reflections in powder XRD patterns that

are broad and diffuse, or completely lacking (Graf and Goldsmith 1956;

Gaines 1977). We define ‘‘disordered dolomite’’ as used here to be a Ca-

Mg carbonate with near dolomite stoichiometry but no evidence of cation

ordering and therefore belonging to the R3c space group. Disordered

dolomite has been termed ‘‘very high-magnesium calcite’’ by Gregg et al.

(2015).

Fluctuation of precipitation conditions, such as pH, and/or a sequence of

solvent-mediated processes, facilitates the formation of disordered

dolomite (Libermann 1967; Deelman 1999). Syntheses of Ca-Mg

carbonates in anhydrous solvent show that Mg/Ca partitioning coefficient

in carbonate crystallized from an anhydrous solvent is about an order of

magnitude higher than that crystallized from an aqueous solution (Xu et al.

2013). Studies point out the positive correlation between the presence of

microbial activity and sedimentary dolomite formation (Hardie 1987;

Compton 1988; Baltzer et al. 1994; Mazzullo et al. 1995; Vasconcelos and

McKenzie 1997; Warren 2000; Van Lith et al. 2002, 2003; Wright and

Wacey 2005; Zhang et al. 2015). Another major aspect of the dolomite

problem pertains to the apparent difficulty with magnesium dehydration

(Xu et al. 2013; Zhang et al. 2013; Gregg et al. 2015; Shen et al. 2015;

Kaczmarek and Thornton 2017). The affinity of water and magnesium

create an energy barrier for the formation of dolomite so that it is hard to

form in a low-temperature setting (Lippmann 1973; Berner 1975; Reddy

and Nancollas 1976; Reddy 1977; Mucci and Morse 1983; Oomori and

Kitano 1987; Falini et al. 1996; Davis et al. 2000; Raz et al. 2000; de

Leeuw and Parker 2001; Higgins and Hu 2005; Stephenson et al. 2008;

Astilleros et al. 2010). Recent studies suggested that dolomitization

requires the presence of dissolved sulfide or polysaccharides as a catalyst

Published Online: June 2018Copyright � 2018, SEPM (Society for Sedimentary Geology) 1527-1404/18/088-679/$03.00

since dissolved sulfide or polysaccharides could be adsorbed onto the

crystal (104) surface, weaken the bond between Mg2þ and surface water,

and thereby lower the energy threshold to form dolomite (Zhang et al.

2012a, 2012b; Shen et al. 2014, 2015; Zhang et al. 2015).

In Mesozoic, Paleozoic, and Proterozoic carbonate sedimentary

deposits, alternating dolomitic and calcitic layers or beds, such as ribbon

rock and stromatolites, are found globally (Sander 1936; Sando 1957;

Fischer 1964; Aitken 1967; Laporte 1967; Matter 1967; Gebelein and

Hoffman 1973; Wanless 1979; Read 1980; Demicco 1983; Knoll and

Swett 1990; Li et al. 2016). The presence of mudcracks, ripple marks, and

bioturbation indicate that these alternating layers were formed in an

intertidal environment (Aitken 1967; Gebelein and Hoffman 1973;

Demicco 1983). Similar textures are observed in Holocene sedimentary

rocks in subtropical and tropical shallow environments ( Logan 1961;

Neumann et al. 1969; Shinn et al. 1969; Gebelein 1972; Gebelein and

Hoffman 1973) and in rare cases found in hypersaline lakes with high

microbial activity without significant grazing (Last et al. 2010, 2012). Our

study is to provide connection between the microbial biomass and dolomite

in the Ordovician rock record. This model of exopolymeric-substances

(EPS)-catalyzed dolomitization explains why a much wider range of

sedimentary dolomite formation on interbedded limestone–dolomite is

common in the ancient sedimentary record (Sando 1957; Aitken 1967;

Matter 1967; Gebelein and Hoffman 1973; Wanless 1979; Read 1980;

Demicco 1983; Hoffman et al. 1998; Hoffman 2011; Liu et al. 2014).

SAMPLES AND GEOLOGIC SETTING

The studied sample was collected from the Middle–Lower Ordovician

St. Paul Group (Row Park and New Market Formation) from the

Massanutten Synclinorium, located in the Central Appalachians in

southern Pennsylvania (Mitchell 1985). The St. Paul Group is overlain

by the Chambersburg Formation, which grades upward into the

Martinsburg Formation. The sample was formed in the upper supratidal

to intertidal setting with presence of mudcracks (Hardie and Ginsburg

1977; Mitchell 1985). With evidence of a shallow environment and no

obvious signs of bioturbation, the sample was assumed to be deposited in

an environment colonized by microbial mats (Mitchell 1985). Modern

analogues of such environments include Andros Island tidal flats and the

Great Bahama Bank (Mitchell 1985). The thin laminae are composed of

distinctive black laminae and gray laminae with average thickness around

0.2 mm (Fig. 1 A, B). Rhombohedral dolomite crystals are exclusively

concentrated within the dark laminae. Mudcracks, scour depressions, and

cryptomicrobial laminates are present in the sample. Similar to this

Ordovician carbonate, Cambrian ribbon rocks are composed of two sets of

beds. Dolomite-dominated beds are yellowish and tan in color due to

oxidation of trace Fe(II) in dolomite layers, while grayish beds with relief

are dominated by calcite. The major difference between Ordovician

carbonate and Cambrian ribbon rock are layer thickness. While the

Ordovician rock has sub-millimeter average thickness, the Cambrian

ribbon rock has thicknesses on the centimeter scale. Mudcracks cut straight

through light layers but are distorted in dark layers (Demicco 1983).

METHODS

XRD analyses were carried out using a Rigaku Rapid II X-ray

diffraction system (Mo Ka radiation). Samples were cut made into thin

chips ~ 100 lm or less thick and glued onto glass capillaries. A 2-D

image-plate detector was used for collecting the diffraction data. The 2-D

images were integrated through Rigaku’s 2DP software and formulated into

traditional 2h vs. intensity patterns using the Jade 9 program. Mineral

percentage and dolomite unit-cell parameters were obtained from Rietveld

refinement using the Jade 9 program.

Average chemical compositions of the dolomite crystals and calcite

microcrystals were determined by electron-probe microanalysis (EPMA).

Elemental data were collected using a CAMECA SX51 instrument with

wavelength-dispersive spectrometers at 15 kV accelerating voltage, 10 nA

beam current, and ~ 3 lm3 interaction volume. A characterized dolomite

standard from Delight, Baltimore (49.43 mol% CaCO3, 50.48 mol%

MgCO3, and 0.09 mol% FeCO3) as used for the EPMA analyses (Zhang

2010). Scanning-electron-microscope (SEM) observations were conducted

using a Hitachi S3400.

A spherical aberration-corrected field-emission gun-scanning transmis-

sion electron microscope (FEG-STEM) (Titan 80-200) operating at 200 kV

at the University of Wisconsin–Madison was used to examine the

microstructures and interface structure of the dolomite-dominated and

magnesian-calcite-dominated lamellae. This instrument has the ability to

image single atoms with ~ 0.08 nm spatial resolution in STEM mode.

Probe current was set at 24.5 pA, and Z-contrast images were collected

with a high-angle annular dark field (HAADF) detector at angle ranges

from 54 to 270 mrad (corresponding to 7.5 (1/A) to 38.2 (1/A) in

reciprocal space). The STEM can be viewed as an inverted conventional

TEM (CTEM) (Kirkland 1998). In STEM, the bright-field (BF) detector

usually collects over a small disc of low-angle coherently scattered

electrons centered on the optical axis of the microscope, while the HAADF

detector collects over an annulus of high-angle incoherently scattered

electrons (Kirkland 1998; Nellist 2007). Intensity of the HAADF image is

strongly related to atomic number (Z) through the Z2 dependence of the

Rutherford scattering cross section. The Z-contrast image, an atomic

resolution HAADF image, can avoid multiple diffractions that commonly

occur in HRTEM and electron-diffraction modes that use elastic coherent

electrons. The BF image is expected to contain diffraction and strain

contrast that is not obvious in HAADF and Z-contrast images (Xu et al.

2014).

For near-infrared spectrum, short-wave infrared (SWIR) imaging was

obtained from a sisuCHEMA imaging station with a hyperspectral camera

(Specim Oy, Oulu, Finland). The light source was a line of quartz-halogen

lamps with a wavelength ranging from 900 nm to 2500 nm. The near-

infrared spectrum collected has both absorbance mode and reflectance

mode. Images were processed using hyperspectrum software developed by

Middleton Spectral Vision, Wisconsin, USA.

Laser-induced-fluorescence (LIF) images are acquired from a macro-

Phor Hyperspectral Fluorescence Imaging System (Middleton Spectral

Vision). The excitation source of the LIF was a 488-nm-wavelength laser.

An M4 Tornado micro-XRF system (Bruker Nano Analytics, Berlin,

Germany) was used to measure the spatial distribution of elements in

submillimeter scale. The measurements were done with a 20 lm spot size

using a molybdenum X-ray source.

RESULTS

Light and dark laminae in the sample examined are prominent under

plane-polarized light (Fig. 2). Horizontal cracks parallel to bedding can be

observed. Most of the euhedral dolomite crystals are concentrated in the

dark layers. Euhedral dolomite crystals, with diameters less than 100 lm,

are surrounded by dark micritic calcite matrix. In contrast, dolomite

crystals are rare, anhedral, and much smaller in the light layers. Many

dolomite crystals are found with microcrystalline calcite inclusions. Small

euhedral pyrite crystals are also occasionally present in the dark laminae.

Some of the euhedral pyrite crystals show subtle crystal zoning, which is a

feature frequently observed in Paleoproterozoic stromatolites (Zentmyer et

al. 2011). While most laminae are planar, some dark laminae are crinkled

and distorted by folding that is interpreted to have occurred during early

compaction; such behavior is consistent with a higher organic concentra-

tion in these layers during deposition and compaction (Sone and Zoback

2013).

Y. FANG AND H. XU680 J S R

Through 42 sets of X-ray diffraction experiments, the sample shows

consistent patterns (Fig. 3). Calcite is identified using ‘‘a’’ reflections

including but not limited to (104), (110), and (006). Comparing with

calcite (space groups R3c), dolomite (space group R3) has a series of extra

diffraction peaks, ‘‘b’’ reflections that violate c glide in calcite R3c. In order

to quantify mineral assemblages within each lamina, the Rietveld

refinement method (MDI Jade 9.0 software) was applied for identified

phases based on fitted XRD profiles. Consistent with the euhedral crystal

distribution from optical images, dolomite is more abundant in the dark

laminae and calcite is more abundant in the light laminae. Dolomite

content in the dark layers is concentrated in ranges from 19.1% to 44.5%

and decrease to less than 15% in most of the light layers. (Table 1, Fig. 4).

FIG. 1.—T-3024 hand sample and ribbon rock from outcrop. A) Back view of T-3024 hand specimen has very thin laminae and trough crossbedding. B) Front view of the

Ordovician carbonate hand sample in Part A. Mudcrack is preserved on the right side of the rock. C) Outcrop image of ribbon rock from Upper Cambrian Conococheague

Formation, Central Appalachians with a US 25¢ coin for scale.

MICROBIAL IMPACT ON SEDIMENTARY DOLOMITE FORMATION FROM A LAMINATED CARBONATEJ S R 681

Other than dolomite enrichment in dark layers, detrital minerals such as

quartz, orthoclase, and trace amounts of chlorite are also enriched in dark

laminae with an average of ~ 29.3% detrital minerals, while light layers

only contain ~ 16.7% detrital minerals on average. Appearance of the

(015) reflection suggests that dolomite in this sample is ordered. The

position of the (104) reflection for dolomite at 2.897(4) A in the dark layers

corresponded to 45 6 1.9% MgCO3 (Reeder and Sheppard 1984; Zhang et

al. 2010).

Scanning electron microscope (SEM) reveals that dark laminae with

cracks cutting through the dark layers are more porous than the adjacent

light layers. Euhedral dolomite in the micrite matrix is visible in

backscattered electron (BSE) images (Fig. 5). BSE images show that the

darker dolomite-rich laminae (~ 0.2 mm) are significantly thinner than

lighter calcite-dominated layers (~ 1.0 mm). BSE imagery and X-ray

energy-dispersive spectra (EDS) indicate that dark layers are composed

mainly of dolomite with calcite inclusions, quartz, orthoclase, pyrite,

chlorite, rutile, and apatite. Results of electron-probe microanalysis

(EPMA) indicate that the average composition of dolomite and calcite

are Ca1.13(Mg0.85Fe0.02)(CO3)2 and Ca1.96(Mg0.04)(CO3)2, respectively.

MgCO3 percentages in dolomite range from 38.5% to 47%, with a mode

around 44% to 45.5%. EPMA further indicates different chemical

composition in the rim and core with an average composition of

Ca1.09(Mg0.90Fe0.01)(CO3)2 in cores and Ca1.12(Mg0.82Fe0.0.06)(CO3)2 in

rims (Table 2, Fig. 6). This variation in Fe content is visible in BSE images

of the dolomite crystals (Fig. 5F).

An electron-diffraction pattern (inserted in the lower-left corner of Fig.

7A) obtained with the scanning transmission electron microscopic (STEM)

image shows strong ‘‘b’’ reflections indicating that dolomite crystals are

well ordered (Fig. 7C). Appearance of (003), (101), (10), and (105)

diffraction spots indicate an ordered dolomite with alternating Ca2þ and

Mg2þ cation layers. Bright-field images show nano-precipitates in dolomite

crystal along ~ (001) and (104) traces (Fig. 7A). Nano-precipitates are

nanometer-scale features precipitated in a certain direction out from the

Ca-rich dolomite matrix. The nano-precipitates are Mg-calcite nano-

precipitates in the dolomite matrix, which is a metastable product during

ordering the Ca-rich dolomite (Shen et al. 2013, 2014). Shen et al. (2013)

has shown that Ca-rich lamellae are parallel to (104) or (110) planes in

order to reduce interfacial strain with host dolomite. This phenomenon has

been observed in Mesoproterozoic (Shen et al. 2014) and Ordovician (Shen

et al. 2013) dolomite. Similar metastable nano-precipitates also occur in

orthopyroxenes as Guinier-Preston zones (Champness and Lorimer 1973;

Xu et al. 2014). Strain contrast in the bright-field image is resulted from

the interface between the precipitate and the dolomite matrix.

Profiling in a Z-contrast image along the (102) trace shows that the

matrix area has periodic Ca-Mg distribution, which suggests that it is

ordered dolomite, while the precipitates are dominated by consecutive Ca

peaks, which shows the precipitates are Ca-rich with decreased cation

ordering (Fig. 7B). Similar magnesian-calcite nano-precipitates were

observed in Ordovician dolomites from the Platteville Formation, western

Wisconsin (Shen et al. 2013). Formation of these magnesian-calcite

precipitates could achieve a lower energy state by removing extra Ca2þ

from the dolomite matrix (Shen et al. 2014). Based on selected area fast-

Fourier transform (FFT) patterns, the matrix area shows clear ‘‘b’’

reflections that indicate high ordering, whereas along the precipitates

‘‘b’’ reflection peaks such as (003) and (105) are either absent or very

weak. The presence of the magnesian-calcite nano-precipitates explains a

phenomenon of ‘‘two dolomite phases’’ in sedimentary dolomite observed

by Drits et al. (2005) where excess-Ca dolomite typically has two phases

with different CaCO3%.

Micro-XRF data, especially Ca and Mg distribution, confirm the

oscillatory distribution of dolomite-dominated and calcite-dominated

laminae (Fig. 8). Ti, Al, and Si signatures come from detrital minerals

such as quartz, orthoclase, rutile, and hydrated silicates such as chlorite.

All of these detrital minerals are concentrated within the dolomite-

dominated laminae, as indicated by correlations among the elemental-

distribution maps. Fe likely substitutes for Mg in dolomite as Fe

distribution overlapped with Mg (Beran and Zemann 1977; Navrotsky et

FIG. 2.—Optical images of sample T-3024. A) Plane-polarized-light and B) cross-polarized-light images of the Ordovician carbonate. Planar laminae with alternating layers

between planar (euhedral) dolomite-dominated layers and microcrystalline calcite layers. Labels in yellow indicate dolomite-dominated layers (labeled D) and calcite-

dominated layers (labeled C).

Y. FANG AND H. XU682 J S R

al. 1999). Fe also is concentrated in pyrite, and the S signature likely results

from the presence of both pyrite and organic materials such as kerogen.

Short-wave infrared (SWIR) spectra show clear dark and light laminae

(Fig. 9A). The SWIR spectrum is presented in reflectance mode in

accordance with previous literature (Gaffey 1987; Morris et al. 2011). In

reflectance mode, spectrum damping indicates absorbance of near-infrared

by the sample (Fig. 9B). In the case for this particular Ordovician

carbonate sample, the absorbance damps have a wavelength around 1947

nm, and 2325 nm for dark layers and 2335 nm for light layers (Fig. 9C).

These absorbance wavelengths indicate the appearance of OH and

carbonate groups (Gaffey 1987; Morris et al. 2011). Accordingly, the

dark layers remain dark in the reflectance mode because darker layers

FIG. 3.—XRD patterns for dark layers and light

layers of Ordovician carbonate using Mo-Kavalue. D, dolomite; C, calcite; Or, orthoclase; Ch,

chlorite. Presence of (01) peak means that

dolomite in this sample is ordered. Dolomite is

scarce or even absent in light layers.

MICROBIAL IMPACT ON SEDIMENTARY DOLOMITE FORMATION FROM A LAMINATED CARBONATEJ S R 683

likely contain abundant hydroxylated organic matter. The 2325 nm

absorbance in dark layers and 2335 nm absorbance in light layers

correspond to characteristic absorbance wavelengths of dolomite and

calcite, respectively (Gaffey 1987). Moreover, using hyperspectral

imaging, the variable spatial concentration of OH and CO32– can be

shown that using 1960 nm and 2320 nm wavelength images in absorbance

mode as reflectance mode will contain background bias. The spatial

distribution of OH is highly correlated with dolomite, based on near-

infrared images using 1960 nm and 2320 nm (Fig. 9D, E).

Based on a previous study, LIF likely detects organic materials in solid

samples (Milori et al. 2006). LIF data on the Ordovician sample show the

same oscillating pattern as do other methods. Intensity of fluorescence in

dark laminae is two to three times higher than fluorescence intensity in

light laminae. Appearance of a fairly broad peak across 500 nm to 750 nm

with a maximum around 550 nm indicates the presence of organic

remnants in the sample (Cecchi et al. 2000) (Fig. 10). The sharp, narrow

peaks at around 650 nm and scattered peaks from 575 nm to 650 nm are

fluorescence due to superimposed dolomite-crystal signals on organic-

material signals (Cecchi et al. 2000). Fluorescence results show that

dolomite and organic matter are highly correlated and concentrated within

the same layers.

DISCUSSION

Understanding the mechanisms and environments of dolomite formation

has been challenging due to the difficulty of abiotically synthesizing

dolomite in the laboratory at low temperature (Land 1998; Gregg et al.

2015) and scarcity of modern examples (Hardie 1987; Warren 2000;

Machel 2004). The struggle to synthesize dolomite, the inability to explain

the formation mechanism of dolomite, and the discrepancy between

abundant ancient record and limited modern dolomite has been referred to

as the ‘‘dolomite problem’’ (Zengler et al. 1980; Machel and Mountjoy

1986; Hardie 1987; Burns et al. 2000; Mazzullo 2000; Warren 2000).

Various hypotheses have been proposed to answer the ‘‘dolomite problem’’

(Zengler et al. 1980; Hardie 1987; Mazzullo 2000; Warren 2000;

Kaczmarek et al. 2017).

The structural differences between dolomite (R3) and calcite (R3c) come

from alternating layers of Ca2þ and Mg2þ and the resulting lower symmetry

in ordered dolomite. Calcite (R3c) and stoichiometric dolomite (R3)

compose a continuous solid solution series with increasing MgCO3 with

Mg-bearing calcite (R3c) and disordered dolomite (R3c) as metastable

phases (Gregg et al. 2015). From previous synthesizing experiments,

aragonite is the kinetically preferred product from a high Mg/Ca solution at

elevated temperature (Kitano 1962; Rushdi et al. 1992; Morse et al. 1997).

Therefore it is believed that Mg2þ ions absorbed on crystal surfaces will

inhibit growth of Ca-Mg carbonate and shift the system towards aragonite

formation (Lippmann 1973; Berner 1975; Reddy and Nancollas 1976;

Reddy 1977; Mucci and Morse 1983; Oomori and Kitano 1987; Falini et

al. 1996; Davis et al. 2000; Raz et al. 2000; de Leeuw and Parker 2001;

Higgins and Hu 2005; Stephenson et al. 2008; Astilleros et al. 2010).

Previous studies suggest that the strong affinity of Mg2þ for water and the

hydration sphere around Mg2þ inhibits Mg2þ from entering the crystal

lattice and therefore impedes growth of Mg-bearing calcite and dolomite

(Lippmann 1973; Mucci and Morse 1983; Davis et al. 2000; Raz et al.

2000; de Leeuw and Parker 2001; Higgins and Hu 2005; Stephenson et al.

2008; Astilleros et al. 2010; Zhang et al. 2012b). More recent anhydrous

solvent synthesis shows that hydrated Mg2þ on crystal growth surfaces,

which prohibits the incorporation of CO32– into the carbonates, is the

actual hindering step rather than the inability of hydrated Mg2þ to enter the

crystal lattice (Xu et al. 2013). Therefore, dehydration of Mg2þ ions

residing on crystal surfaces is the key step in dolomite formation.

On the other hand, polysaccharides, a major component of extracellular

polymeric substances (EPS), likely promote the growth of disordered

TABLE 1.—Weight percentages of minerals for individual dark or light

laminae from Rietveld analyses of the XRD results. L stands for light

layers and D stands for dark layers.

Dolomite Calcite Quartz Orthoclase Total

1L 9.6 71.2 1.7 17.5 100.0

2D 38.8 36.2 2.8 22.2 100.0

3L 4.0 78.8 3.4 13.8 100.0

4D 47.2 35.8 2.2 14.8 100.0

5L 19.0 71.6 1.1 8.3 100.0

6D 45.9 34.6 3.1 16.4 100.0

7D 8.3 40.5 8.2 43.0 100.0

8L 10.1 76.7 1.7 11.4 99.9

9L 3.5 78.8 3.5 13.9 99.7

10D 41.2 35.9 3.8 19.1 100.0

11L 2.8 85.7 1.6 9.8 99.9

12D 31.4 31.7 2.5 34.4 100.0

13L 2.7 83.2 2.2 11.8 99.9

14D 34.3 16.0 1.0 48.7 100.0

15L 3.0 81.0 3.0 12.7 99.7

16D 11.4 56.9 1.9 29.9 100.1

17L 2.5 86.0 1.3 10.1 99.9

18D 23.3 50.1 3.1 23.4 99.9

19D 38.4 36.3 2.1 23.2 100.0

20L 14.9 71.3 0.8 12.9 99.9

21D 17.4 45.1 5.0 32.4 99.9

22L 2.2 73.2 5.8 18.8 100.0

23D 40.4 23.5 2.7 33.4 100.0

24D 40.2 44.1 1.8 13.9 100.0

25D 39.3 37.1 3.0 20.6 100.0

26L 17.9 62.2 1.6 18.3 100.0

27D 22.3 47.6 3.1 27.0 100.0

28L 14.7 68.5 1.9 14.9 100.0

29D 31.6 52.2 2.3 13.9 100.0

30D 30.8 37.0 3.2 29.0 100.0

31L 4.0 65.2 5.8 25.0 100.0

32D 50.0 23.8 2.8 23.4 100.0

FIG. 4.—Ternary plot of minerals (wt. %) calculated using Rietveld analysis

method. Dolomite is dominant in the dark layers, whereas calcite is dominant in the

light layers.

Y. FANG AND H. XU684 J S R

FIG. 5.—BSE images of the Ordovician carbonate. A) Overview of multiple layers with clearly defined dolomite-dominated layers and microcrystalline calcite layers. B)

Close-up of layer 1 with euhedral dolomite crystal. Bright spots are mostly pyrite with a small amount of rutile and trace apatite. C) High-magnification view of layer 1 with a

thin horizontal fracture going through it. Minerals are distributed along the fracture with a fuzzy dolomite grain boundary. D) Image of calcite-dominated layer 2 with few

euhedral dolomite crystals in a microcrystalline calcite matrix. Pyrite and rutile are absent in layer 2. E) Part of layer 3 that is composed of two thin layers of euhedral

dolomite. Pyrite, apatite, and a trace amount of mica follow the orientation of the dolomite layers. F) Close-up view of a dolomite crystal. Mineral inclusions and vacancies in

the dolomite are common. Color differences in the rim and core of dolomite results from differences in CaCO3% and FeCO3% of the crystals.

MICROBIAL IMPACT ON SEDIMENTARY DOLOMITE FORMATION FROM A LAMINATED CARBONATEJ S R 685

dolomite (Zhang et al. 2012), promote calcite growth, inhibit aragonite

formation (Kawano and Hwang 2011), and impact calcite morphology

during precipitation (Braissant et al. 2003). Atomic simulation shows that

surface-attached trimannose can effectively lower the dehydration energy

of surface water from its hydrophobic functional group (Shen et al. 2015).

Zhang et al. (2015) used non-metabolizing consortium biomass and bound

EPS, derived from the consortium biomass, as a catalyst and successfully

synthesized disordered dolomite at low temperature. Roberts et al. (2013)

reported bioinorganic synthesis of dolomite at low temperature with

carboxylated polystyrene spheres. However, we reinterpret the results from

Roberts et al. (2013) as showing no sign of dolomite, as we believe that

they misinterpreted the aragonite (002) peak as the dolomite (104) in their

XRD pattern, the aragonite [3 1 14] zone-axis for the dolomite [2 1 1]

zone-axis in their electron diffraction pattern, and CaO nano-crystals

resulted from electron-beam-induced decomposition of aragonite as

dolomite in their HRTEM image.

The Ordovician carbonate sample studied here is composed of dolomite-

dominated (dark) and calcite-dominated (light) laminae. Based on optical

images, euhedral (planar) dolomite crystals in the dolomite-dominated

layers suggest that the dolomite formed during early diagenesis (Sibley and

Gregg 1987) with growth temperature lower than 508C (Gregg and Sibley

1984; Warren 2000; Ferry et al. 2011). The chemical zoning in optics and

BSE images are the result of variation of Fe and Mn in the rim and core

with less Fe in the core (Choquette 1980; Fairchild 1983; Gawthorpe 1987;

Warren 2000). EDS and EPMA also show dolomite crystals have excess

CaCO3, but XRD and electron-diffraction patterns obtained from TEM

indicate that the dolomite crystals are well ordered. Z-contrast images from

STEM show that Ca-rich precipitates occur locally in the dolomite matrix

(Fig. 7). Because intensity depends on atomic number and thickness, in a

sample of consistent thickness the change in intensity indicates a change in

atomic number. Therefore, the increase in intensity in some Mg2þ sites

along the [104] direction indicates that extra Ca atoms are substituted into

the Mg2þ positions. These substitutions destroy the Ca2þ-Mg2þ repetition

in ideal dolomite, and result in Ca-rich precipitates (Shen et al. 2013,

2014). An order parameter was proposed to describe the ordering state of

stoichiometric dolomite (Reeder and Wenk 1983; Antao et al. 2004)

s ¼ 2xCa � 1; or s ¼ 2xMg � 1 ð1Þ

where xCa and xMg are the occupancies of Ca at Ca site, and Mg in Mg site.

When xCa ¼ 1, s ¼ 1 complete order dolomite, whereas xCa ¼ 0.5, s ¼ 0

complete disordered dolomite with R3c symmetry. Using Z-contrast image

simulation and the relative intensity of cations along the [104] direction,

TABLE 2.—Electron-probe microanalysis (EPMA) data of T-3024 in oxide percent and normalized values according to dolomite stoichiometry. Mole %

of Cc ¼ CaCO3 %.

Point No.

Oxide % Normalized value

Mole % of CcCaO MgO FeO MnO Total Ca Mg Fe Mn Total

T3-024-1 1 Core 32.167 18.529 0.470 0.000 98.669 1.106 0.887 0.007 0.000 2.000 55.33

2 Rim 31.456 18.592 0.318 0.034 98.13 1.095 0.901 0.004 0.000 2.000 54.74

T3-024-2 3 Core 31.681 19.168 0.597 0.063 98.991 1.081 0.910 0.008 0.001 2.000 54.05

4 Rim 32.047 17.150 2.358 0.080 98.696 1.126 0.837 0.036 0.001 2.000 56.28

5 Core 31.340 19.175 0.673 0.000 98.753 1.076 0.915 0.009 0.000 2.000 53.77

6 Rim 32.762 18.170 0.495 0.017 98.827 1.125 0.868 0.007 0.000 2.000 56.24

T3-024-3 7 Core 31.998 17.858 1.384 0.080 98.626 1.114 0.865 0.020 0.001 2.000 55.69

8 Rim 32.245 18.079 1.749 0.029 99.207 1.109 0.866 0.025 0.000 2.000 55.46

9 Core 31.340 18.174 0.279 0.006 97.645 1.105 0.891 0.004 0.000 2.000 55.23

10 Rim 31.638 18.606 0.724 0.011 98.526 1.095 0.895 0.010 0.000 2.000 54.71

T3-024-4 11 Core 30.493 19.454 0.686 0.000 98.383 1.055 0.936 0.009 0.000 2.000 52.73

12 Rim 32.837 15.765 2.609 0.057 98.229 1.174 0.784 0.041 0.001 2.000 58.71

13 Core 31.344 19.184 0.406 0.000 98.587 1.077 0.917 0.006 0.000 2.000 53.84

14 Rim 31.084 16.492 2.938 0.034 97.774 1.124 0.830 0.045 0.001 2.000 56.23

15 Core 31.819 18.810 0.648 0.023 98.791 1.092 0.899 0.009 0.000 2.000 54.61

16 Rim 33.700 16.747 2.255 0.000 99.428 1.163 0.804 0.033 0.000 2.000 58.13

Average Core Average 31.523 18.794 0.643 0.022 98.556 1.088 0.903 0.009 0.000 2.000 54.49

Rim Average 32.221 17.45 1.681 0.033 98.602 1.127 0.848 0.025 0.000 2.000 56.31

FIG. 6.—Ternary (Ca–Mg–Fe) plot of EPMA

results. Compared to the core composition, the

rim generally has a higher amount of Fe2þ and

less Mg2þ and Ca2þ.

Y. FANG AND H. XU686 J S R

the ordering of these dolomites is shown to be close to s ¼ 0.9 (Fig. 11),

excluding the Ca-rich precipitates.

XRD results indicate that dolomite-dominated layers have significantly

larger quantities of detrital quartz and feldspar than the calcite-dominated

layers (Table 1). Previous studies have recognized the relationship between

microbial organisms and input of detrital minerals. Black (1933) observed

that an environment with periodical flooding would promote growth of

algal mat by providing materials and nutrition from continental weathering.

Baltzer et al. (1994) correlated distributions of dolomite with total organic

carbon in Holocene Ras Ghanada sediments. Hardie and Ginsburg (1977)

FIG. 7.—A) Bright-field (BF), B) Z-contrast

images and their corresponding fast Fourier

transform (FFT) (lower left) along [010] zone axis

of the Ordovician dolomite. Ca-rich precipitates

approximately parallel to (104) and subparallel to

(001). Linear feature in BF image likely come

from strain contrast between precipitates and

matrix. Nano-precipitates are brighter in Z-

contrast image such as the one above the dashed

(104) trace. FFT from precipitate a) shows

diffused and weak ‘‘b’’ reflections that indicate a

decrease in order state in the precipitate, whereas

an area without precipitate b) shows sharp ‘‘b’’

reflections and better resolution. The profile along

the (102) trace shows disorder and excess of Ca

along the precipitate. C) Selected-area electron

diffraction pattern from [010] zone axis of the

Ordovician dolomite showing of distinctive (003),

(101), and (105) ‘‘b’’ reflections, which indicate

that the dolomite is ordered. Ask author to add C

label

MICROBIAL IMPACT ON SEDIMENTARY DOLOMITE FORMATION FROM A LAMINATED CARBONATEJ S R 687

and Mitchell (1985) show that dolomite formed in a storm-influenced tidal

setting. In the sample studied here, flooding may have increased the input

of detrital minerals into the microbial mat, and this could be the reason

larger amounts of quartz and feldspar are present in the dolomite laminae.

Because detrital minerals are enriched in the dolomite laminae (dark layer),

they might form during periods of significantly more frequent flooding

which presumably would be during the summer season, with more sunlight

exposure. Higher temperature, more sunlight, and more nutrients would

have promoted microbial growth and thus provide more organic materials

during deposition, which explains the higher porosity, darker color, and

more plastic behavior of the dolomite-dominated layers. Previous studies

have shown that the storm season can significantly impact the pattern of

carbonate precipitation through change in nutrients and water chemistry

(Folk and Siedlecka 1974; Boero 1996; Petrash et al. 2016). Mudcracks are

highly distorted in the dark layers; this suggests a much larger volume

shrinkage in the dark layers during deposition and later diagenesis than in

the light layers. A larger volume decrease in dark layers indicates higher

organic content in the dolomite dominated layers during their deposition

(Demicco 1983).

In order to spatially correlate organic matter and dolomite in the sample,

the near-infrared method was applied. Based on previous studies, for a

1000–2500 nm wavelength range, maximum absorption would occur at

FIG. 8.—XRF elemental maps of T-3024 thin section. A) Cross-polarized-light image of T-3024 thin section. B–H) Elemental maps of Ca, Mg, Fe, S, Al, Si, and Ti,

respectively.

Y. FANG AND H. XU688 J S R

1960 nm for OH, 2320 nm for dolomite, and 2333 nm for calcite (Gaffey

1987; Morris et al. 2011; Morris et al. 2010). Maximum wavelength for

absorption for this study was slightly different from previous work, with

1947 nm for OH, 2325 nm for dolomite, and 2335 nm for calcite (Fig. 9).

The absorption wavelength was determined by bond distance within the

materials. The difference between this study and the literature might be due

to Ca-rich dolomite (this study) rather than stoichiometric dolomite. The

extra Ca2þ in the crystal lattice will increase the lattice parameters of

dolomite from a¼ 4.8079 A c¼ 16.0100 A (Graf 1961) of ideal dolomite

with 50% MgCO3 to a¼4.829 A c¼16.067A for the Ordovician dolomite

with 45.42% of MgCO3. The increase in lattice parameters gives rise to

increase of bond distance between cation and anion, which affects the

FIG. 9.—SWIR images of the Ordovician carbonate hand sample in A) absorbance mode, B) reflectance mode. C) SWIR spectrum of light laminae (red) and dark laminae

(dark) in reflectance mode. Wavelengths from measurement deviated slightly from referenced 1960 nm for OH, 2320 nm for dolomite, and 2330 nm for calcite absorbance

bands (Gaffey 1987). Images showing relative concentration of D) 1960 nm for OH and E) 2320 nm dolomite across the sample. Spatial distribution of OH and dolomite are

matched from the two images. Difference in measured wavelength is probably due to instrumental difference, and dolomite in this study is slightly Ca-rich.

MICROBIAL IMPACT ON SEDIMENTARY DOLOMITE FORMATION FROM A LAMINATED CARBONATEJ S R 689

FIG. 10.—Laser-induced fluorescence (LIF) image for the Ordovician carbonate. Three different components are identified across the image from deconvolution using

KemoQuant software. The spectrum of the three components Parts B, C, and D correspond with each component’s spatial distribution across the sample in Parts E, F, and G,

respectively. Spectrum in Part B corresponds to fluorescence from dolomite while spectra in Parts C and D correspond to organic remnant.

Y. FANG AND H. XU690 J S R

FIG. 11.—Noise-filtered Z-contrast image and simulated Z-contrast image of Ca-rich dolomite with ordering state of s¼0.9 ordering. The relative intensity from the Z-contrast

image and simulated image along A, B) [001], C, D) [104], E, F) [102�

] directions matched indicate the Ordovician dolomite has approximately state of s¼ 0.9 ordering.

MICROBIAL IMPACT ON SEDIMENTARY DOLOMITE FORMATION FROM A LAMINATED CARBONATEJ S R 691

wavelength of absorption (Gaffey 1987). Figure 9D and E show the

distribution of OH and dolomite based on 1960 nm and 2320 nm

respectively in absorbance mode. It is clear that OH and dolomite are

highly correlated to each other. However, OH signals could come from

either organic or hydrated minerals or both. In order to confirm the organic

distribution rather than the combination of both organic and hydrous

silicates, laser-induced fluorescence was used to produce LIF a signature

fluorescence spectrum for organic matter (Lichtenthaler et al. 1992;

Alberotanza et al. 1995; Tiano et al. 1995; Cecchi et al. 2000; Milori et al.

2006). LIF results from the Ordovician carbonate have identified three

components across the sample. Component 1 (Fig. 10B) is an indication of

dolomite fluorescence, while components 2 (Fig. 10C) and 3 (Fig. 10D) are

signals coming from organic materials (Cecchi et al. 2000). The spatial

distribution of components 1, 2, and 3 (Fig. 10E, F, G) are strongly

correlated in this sub-millimeter-laminated sample. Combining results

from both near-infrared and LIF, the spatial correlation between dolomite

and organic remnants is confirmed. The hypothesis that polysaccharides or

EPS functioning as a catalyst to promote sedimentary dolomite formation

is able to explain the confinement of dolomite and organics to sub-

millimeter fine laminae.

Hardie and Ginsburg (1977) and Mitchell (1985) suggested that the

planar micro-laminated carbonate forms in a sticky biomat in a tidal flat.

The enrichment of organic materials in the dolomite-dominated layer

during deposition likely forms a membrane-like structure from the

polysaccharides in the EPS, which prohibits the exchange of large

molecules but allows small ions and water to pass through. The membrane-

like structures lower the fluidity of the organic materials. The confinements

of membrane-like structures results in the sharply defined micro-laminated

structure in the sample. After deposition, dolomitization would occur

preferentially only in the membrane-like structure that is enriched in

organic materials. With seawater providing Mg2þ into both layers, only

layers with trapped organic materials would have subsequent CO32–

bonded with surface Mg2þ in Mg-calcite and would start to form

disordered and weakly ordered dolomite:

2Ca1�xMgx CO3ð Þ sð Þ þ 2 1� xð ÞMg2þ aqð ÞMg calcite

! CaMg CO3ð Þ2 sð Þ þ x� yð Þ2xCa2þ aqð Þdolomite

; x.y

Therefore, this study suggests that the formation of dolomite layers

started with organic matter that catalyzed dolomitization in the organic

enriched layers around calcite or high-magnesium-calcite seed crystals.

This explains the common calcite micro-inclusions observed in dolomite

crystals in this sample. As sediments continue to be deposited on top of

this organic layer, microorganisms started to decompose, shifting the

system towards a reducing micro-environment (Des Marais 2003).

Increasing burial depth eventually shifts the microbial activity to

fermentation and sulfate reduction (Irwin et al. 1977; Pisciotto et al.

1981; Kelts and McKenzie 1982; Mazzullo 2000). HS– will bind with free

Fe2þ to precipitate pyrite in the organic-rich layers. Studies of Mg isotopes

in ribbon rock shows that the Mg source for dolomite is more likely to

come from contemporaneous seawater than from specific dolomitizing

fluid from Mg released by clay (Li et al. 2016). Therefore, limited

distribution of microbial biomass could explain the periodic distribution in

Ordovician carbonate, as all layers would have been exposed to the same

flux of Mg.

Frequent environmental oscillation was likely responsible for the thin

laminae in the sample under study. The correlation between dolomite

formation and environment could be derived from the relationship between

dolomitization and microbial biomass, because bioactivity is sensitive to

environmental changes (Hardie 1987; Compton 1988; Vasconcelos and

McKenzie 1997; Mazzullo 2000; Warren 2000; Van Lith et al. 2002, 2003;

Wright and Wacey 2005). Organic remnants, which are most likely to be in

the form of kerogen in this Ordovician carbonate, are concentrated within

the dolomite-abundant laminae. SWIR and LIF data show the spatial

correlation between organics and dolomite in submillimeter-scale laminae,

which indicates that the formation of dolomite requires the presence of

microbial EPS as a catalyst. According to Lokier and Steuber (2008), the

average sedimentation rate for carbonate is around 0.29 mm per year. Paull

et al. (1992) have shown that the average sedimentation rate for

stromatolites is about 0.16 mm/yr, which is within the same magnitude

as this Ordovician sample. The Ordovician sample studied here has an

average width of individual laminae of approximately 0.2 mm.

Quantitative analysis on sedimentation rate using average spectral misfit

(ASM) with Monte Carlo spectra simulations (Meyers et al. 2012; Meyers

and Sageman 2007) shows that the sedimentation rate of this sample is

calculated to be 0.098 mm/yr. Because the sedimentation is calculated

using the rock sample, the effect of volume changes from organic

decomposition and compaction during dolomitization is neglected.

Therefore, 0.098 mm/yr might be an underestimation of annual

sedimentation rate. Based on previous studies, the Ordovician carbonate

sample with submillimeter laminae could potentially indicate annual cycles

with a pair of dark and light laminae where dark layers correspond to warm

and nutrient-rich high-energy summer conditions with more organics and

light layers indicate cool, nutrient-deprived, and less productive winter

conditions.

Kendall and Skipwith (1968) found similar features of planar laminae to

those studied here in upper intertidal and supratidal flats in Khor al Bazam,

United Arab Emirates. Ribbon rock from the Conococheague Formation,

Upper Cambrian, shows alternating limestone-dolomite beds that were also

inferred to be formed in a tidal-flat environment (Fig. 1C) (Demicco 1983).

A sample of ribbon rock (see Supplemental Material, File 1) displays

planar laminations similar to the Ordovician carbonates studied here, but

with much thicker beds. XRD patterns and optical examinations

demonstrate that the dark layers in the ribbon rock are enriched in

euhedral dolomite and light layers are dominated by microcrystalline

calcite (Supplemental Material, File 2). Similarity in mineral assemblage

for individual layers and morphology for the ribbon rock and the

Ordovician carbonate implies that dolomites in ribbon rock were likely to

be the product of the changes in climate that lead to different microbial

activity, which influence the amount of microbial EPS. Other factors may

also have influenced carbonate precipitation in these ancient rocks, such as

the climatic effect on the microbial community morphology and the

species of microbes involved. The difference in laminae thickness was

produced by different mechanisms driving environmental oscillation.

Therefore, the Ordovician carbonate is not a unique example of local

sedimentary dolomite precipitation but a case with well-defined constraints

to explain dolomite formation in ambient conditions requiring the presence

of EPS in a microbial biomat as a catalyst.

This model has the potential to address the formation of the ‘‘cap

dolostone’’ observed at the melting of ‘‘Snowball Earth’’ glaciation events.

Paleomagnetic data have shown that a cap dolomite in the Nuccaleena and

Ol formations were presumably formed in a warm, subtropical location

(Kravchinsky et al. 2010; Evans and Raub 2011; Williams et al. 2011; Liu

et al. 2014) with nutrient-rich continental surface runoff and a high water

table. Continental erosion providing nutrients such as life-limiting elements

such as phosphorus promoted microorganism growth at lower latitudes

(Shields 2005; Kunzmann et al. 2013; Liu et al. 2014). Increases in

primary productivity in these locations provided a large quantity of

microbial biomass, including EPS, during deposition and therefore likely

catalyzed the formation of the cap dolostone.

CONCLUSION

The presence of sedimentary dolomite in the studied samples indicates

an environment with thriving microorganisms that could provide catalysts

Y. FANG AND H. XU692 J S R

such as polysaccharides (the main component in microbial EPS). The

strong correlation between the spatial distribution of dolomite and organics

in submillimeter-laminated carbonate rock with sharp boundaries supports

the hypothesis of microbial EPS promoting dolomite growth from ambient

synthesizing experiments and molecular-dynamics modeling work. It is

also likely that heterotrophy of organic matter in sediments produce

suitable conditions for dolomite formation.

Using XRD, SEM, LIF, SWIR, and XRF, the spatial distribution of

mineral phases, elements, and organic content of the well-constrained

Ordovician carbonate with alternating dolomite–calcite layers, were

determined. By applying multiple methods on the same area, spatial

correlation between dolomite and organic remnants was established. The

evidence presented here supports the hypothesis that dolomite formation

requires the presence of organic catalysis such as microbial EPS. Thus, the

occurrence of similar sedimentary dolomite in the rock record implies an

environment suitable to supporting a relatively high degree of microbial

activity. Similar dolomitization patterns are observed in Cambrian ribbon

rocks and in the Neoproterozoic ‘‘cap dolostone,’’ indicating that this is not

an isolated occurrence of dolomitization associated with such microbial

activity. The differences in thickness of laminae between Ordovician

carbonate, ribbon rock, and ‘‘cap dolostone’’ were probably not the result

of different dolomitization mechanisms but the stability of a suitable

environment for microorganisms to thrive and the periodicity of

paleoenvironmental changes.

Oscillatory appearance in other types of sedimentary carbonate records

could also then be used to interpret paleoenvironmental changes. Changes

in climate may have resulted in variation in microbial activity, which

determined the amount of organic matter such as polysaccharide and EPS,

available to catalyze dolomite growth. Other factors such as changes in

microbial community structures might impact the amount of catalysts for

dolomite formation, but the exact contribution of these other factors are yet

to be determined.

SUPPLEMENTAL MATERIALS

Supplemental Files 1 and 2 are available from JSR’s Data Archive:

https://www.sepm.org/pages.aspx?pageid=229.

ACKNOWLEDGMENTS

We thank Drs. Jay Gregg, Stephen Kaczmarek, John Southard, and an

anonymous reviewer for their constructive suggestions and comments, Profs.

John Valley and Phil Brown for their helpful advice, Dr. Zhizhang Shen for

helping with the initial XRD experiments, Dr. Hiromi Konishi for the STEM

work, Drs. Gaber Kemeny and Chris Draves at Spectral Vision Middleton for

SWIR and LIF data collection, and Bruker AXS Inc. for XRF experiments. This

work is supported by NASA Astrobiology Institute (NNA 13AA94A).

REFERENCES

AITKEN, J.D., 1967, Classification and environmental significance of cryptalgal limestones

and dolomites, wiht illustrations from the Cambrian and Ordovician of southwestern

Alberta: Journal of Sedimentary Petrology, v. 37, p. 1163–1178.

ALBEROTANZA, L., COVA, P.L., RAMASCO, C., VIANELLO, S., BAZZANI, M., CECCHI, G., PANTANI,

L., RAIMONDI, V., RAGNARSON, P., SVANBERG, S., AND WALLINDER, E., 1995, Yellow

substance and chlorophyll measurements in the Venice lagoon using laser-induced

fluorescence: Advances in Remote Sensing, v. 3, p. 102–111.

ANTAO, S.M., MULDER, W.H., HASSAN, I., CRICHTON, W.A., AND PARISE, J.B., 2004, Cation

disorder in dolomite, CaMg(CO3)2, and its influence on the aragonite þ magnesite �dolomite dolomite reaction boundary: American Mineralogist, v. 89, p. 1142–1147.

ASTILLEROS, J.M., FERNANDEZ-DIAZ, L., AND PUTNIS, A., 2010, The role of magnesium in the

growth of calcite: an AFM study: Chemical Geology, v. 271, p. 52–58.

BALTZER, F., KENIG, F., BOICHARD, R., PLAZIAT, J.C., AND PURSER, B.H., 1994, Organic matter

distribution, water circulation and dolomitization beneath the Abu Dhabi sabkha (United

Arab Emirates), in Purser, B., Tucker, M., and Zenger, D., eds., Dolomites, A Volume in

Honour of Dolomieu: International Association of Sedimentologists, Special Publication

21, p. 409–427.

BERAN, A., AND ZEMANN, J., 1977, Refinement and comparison of the crystal structures of a

dolomite and of an Fe-rich ankerite: Tschermaks Mineralogische und Petrographische

Mitteilungen, v. 24, p. 279–286.

BERESKIN, S.R., MORGAN, C.D., AND MCCLURE, K.P., 2004, Descriptions, Petrology,

Photographs, and Photomicrographs of Core from the Green River Formation, South-

Central Uinta Basin, Utah (MP-04-2): Utah Geological Survey, Miscellaneous

Publication, 208 p.

BERNER, R.A., 1975, The role of magnesium in the crystal growth of calcite and aragonite

from sea water: Geochimica et Cosmochimica Acta, v. 39, p. 489–504.

BLACK, M., 1933, The Algal sediments of Andros Island, Bahamas: The Royal Society of

London, Philosophical Transactions, Series B, Containing Papers of a Biological

Character, v. 222, p. 165–192.

BOERO, F., 1996, Episodic events: the relevance to ecology and evolution: Marine Ecology,

v. 17, p. 237–250.

BRAISSANT, O., CAILLEAU, G., DUPRAZ, C., AND VERRECCHIA, E.P., 2003, Bacterially induced

mineralization of calcium carbonate in terrestrial environments: the role of exopoly-

saccharides and amino acids: Journal of Sedimentary Research, v. 73, p. 485–490.

BRAITHWAITE, C.J.R., RIZZI, G., AND DARKE, G., 2004, The geometry and petrogenesis of

dolomite hydrocarbon reservoirs: Introduction, in Braithwaite, C.J.R., Rizzi, G., and

Darke, G., eds., The Geometry and Petrogenesis of Dolomite Hydrocarbon Reserviors:

Geological Society of London, Special Publication 235, p. 1–6.

BURNS, S.J., MCKENZIE, J.A., AND VASCONCELOS, C., 2000, Dolomite formation and

biogeochemical cycles in the Phanerozoic: Sedimentology, v. 47, p. 49–61.

CECCHI, G., PANTANI, L., RAIMONDI, V., TOMASELLI, L., LAMENTI, G., TIANO, P., AND CHIARI,

R., 2000, Fluorescence lidar technique for the remote sensing of stone monuments:

Journal of Cultural Heritage, v. 1, p. 29–36.

CHAMPNESS, P.E., AND LORIMER, G.W., 1973, Precipitation (exosolution) in an orthopyrox-

ene: Journal of Material Science, v. 8, p. 467–474.

CHOQUETTE, P.W., 1980, Mississippian non-supratidal dolomite, Ste. Genevieve Limestone,

Illinois Basin: evidence for mixed-water dolomitization, in Zenger, D.H., Dunham, J.B.,

and Ethington, R.L., eds., Concepts and Models of Dolomitization: SEPM, Special

Publication 28, p. 163–196.

COMPTON, J.S., 1988, Degree of supersaturation and precipitation of organogenic dolomite:

Geology, v. 16, p. 318–321.

DAVIS, K.., DOVE, P.M., AND DE YEREO, J.J., 2000, The role of Mg2þ as an impurity in calcite

growth: Science, v. 290, p. 1134–1137.

DEELMAN, B.J.C., 1999, Low-temperature nucleation of magnesite and dolomite: Neues

Jahrbuch Fur Mineralogie, Monatshefte, v. 7, p. 289–302.

DE LEEUW, N., AND PARKER, S., 2001, Surface-water interactions in the dolomite problem:

Physical Chemistry Chemical Physics, v. 3, p. 3217–3221.

DEMICCO, R.V., 1983, Wavy and lenticular-bedded carboante ribbon rocks of the Upper

Cambrian Conococheague limestone, Central Appalachians: Journal of Sedimentary

Petrology, v. 53, p. 1121–1132.

DES MARAIS, D.J., 2003, Biogeochemistry of hypersaline microbial mats illustrates the

dynamics of modern microbial ecosystems and the early evolution of the biosphere:

Biological Bulletin, v. 204, p. 160–167.

DRITS, V.A., MCCARTY, D.K., SAKHAROV, B., AND MILLIKEN, K.L., 2005, New insight into

structural and compositional variability in some ancient excess-Ca dolomite: The

Canadian Mineralogist, v. 43, p. 1255–1290.

EVANS, D.A.D., AND RAUB, T.D., 2011, Neoproterozoic glacial palaeolatitudes: a global

update, in Arnaud, E., Halverson, G.P., and Sheilds-Zhou, G., eds., The Geological

Record of Neoproterozoic Glaciations: Geological Society of London, Memoir 36, p.

93–112.

FAIRCHILD, I., 1983, Chemical controls of cathodoluminescence of natural dolomites and

calcites: new data and review: Sedimentology, v. 30, p. 579–583.

FALINI, G., GAZZANOB, M., AND RIPAMONTI, A., 1996, Magnesium calcite crystallization from

water–alcohol mixtures: Chemical Communications, v. 9, p. 1037–1038.

FERRY, J.M., PASSEY, B.H., VASCONCELOS, C., AND EILER, J.M., 2011, Formation of dolomite

at 40–80 8C in the Latemar carbonate buildup, Dolomites, Italy, from clumped isotope

thermometry: Geology, v. 39, p. 571–574.

FISCHER, A.G., 1964, The Lofer cyclothems of the alpine Triassic: Kansas Geological

Survey, Bulletin, v. 169, p. 107–149.

FOLK, R.L., AND SIEDLECKA, A., 1974, The ‘‘schizohaline’’ environment: its sedimentary and

diagenetic fabrics as exemplified by late Paleozoic rocks of Bear Island, Svalbard:

Sedimentary Geology, v. 11, p. 1–15.

GAFFEY, S., 1987, Spectral reflectance of carbonate minerals in the visible and near infrared

(0.35–2.55 um): anhydrous carbonate minerals: Journal of Geophysical Research, v. 92,

p. 1429–1440.

GAINES, A.M., 1977, Protodolomite redefined: Journal of Sedimentary Petrology, v. 47, p.

543–546.

GAWTHORPE, R.L., 1987, Burial dolomitization and porosity development in a mixed

carbonate–clastic sequence: an example from the Bowland Basin, northern England:

Sedimentology, v. 34, p. 533–558.

GEBELEIN, C.D., 1972, Sedimentology and ecology of a recent carbonate facies mosaic,

Cape Sable, Florida [unpublished Ph.D. dissertation]: Brown University, 237 p.

GEBELEIN, C.D., AND HOFFMAN, P., 1973, Algal origin of dolomite laminations in

stromatolitic limestone: Journal of Sedimentary Petrology, v. 43, p. 603–613.

GOLDSMITH, J.R., AND GRAF, D.L., 1958, Strctural and compositional variations in some

natural dolomites: The Journal of Geology, v. 66, p. 678–693.

MICROBIAL IMPACT ON SEDIMENTARY DOLOMITE FORMATION FROM A LAMINATED CARBONATEJ S R 693

GRAF, D.L., 1961, Crystallographic tables for the rhombohedral carbonates: American

Mineralogist, v. 46, p. 1283–1316.

GREGG, J.M., AND SIBLEY, D.F., 1984, Epigenetic dolomitization: the origin of xenotopic

dolomite texture: Journal of Sedimentary Petrology, v. 54, p. 908–931.

GREGG, J.M., BISH, D.L., KACZMAREK, S.E., AND MACHEL, H.G., 2015, Mineralogy,

nucleation and growth of dolomite in the laboratory and sedimentary environment: a

review: Sedimentology, v. 62, p. 1749–1769.

HARDIE, L.A., 1987, Dolomitization; a critical view of some current views: Journal of

Sedimentary Research, v. 57, p. 166–183.

HARDIE, L.A., AND GINSBURG, R.N., 1977, Layering: the origin and environmental

significance of lamination and thin bedding, in Hardie, L.A., ed., Sedimentation on the

Modern Carbonate Tidal Flats of Northwest Andros Island, Bahamas: The Johns

Hopkins University, Studies in Geology, v. 22, p. 50–123.

HIGGINS, S.R., AND HU, X., 2005, Self-limiting growth on dolomite: experimental

observations with in situ atomic force microscopy: Geochimica et Cosmochimica Acta,

v. 69, p. 2085–2094.

HOFFMAN, P.F., 2011, Strange bedfellows: glacial diamictite and cap carbonate from the

Marinoan (635 Ma) glaciation in Namibia: Sedimentology, v. 58, p. 57–119.

HOFFMAN, P.F., KAUFMAN, A.J., HALVERSON, G.P., AND SCHRAG, D.P., 1998, A Neoproterozoic

Snowball Earth: Science, v. 281, p. 1342–1346.

IRWIN, H., CURTIS, C., AND COLEMAN, M., 1977, Isotopic evidence for source of diagenetic

carbonates formed during burial of organic-rich sediments: Nature, v. 269, p. 209–213.

KACZMAREK, S.E., AND THORNTON, B.P., 2017, The effect of temperature on stoichiometry,

cation ordering, and reaction rate in high-temperature dolomitization experiments:

Chemical Geology, v. 468, p. 32–41.

KACZMAREK, S.E., GREGG, J.M., BISH, D., MACHEL, H., AND FOUKE, B., 2017, Dolomite,

very-high magnesium calcite, and microbes: implications for the microbial model of

dolomitization, in MacNeil, A., Lonnee, J., and Wood, R., eds., Characterization and

Modeling of Carbonates, Mountjoy Symposium 1: SEPM, Special Publication 109, 17 p.

KAWANO, M., AND HWANG, J., 2011, Roles of microbial acidic polysaccharides in

precipitation rate and polymorph of calcium carbonate minerals: Applied Clay Science,

v. 51, p. 484–490.

KELTS, K., AND MCKENZIE, J.A., 1982, Diagenetic dolomite formation in quaternary anoxic

diatomaceous muds of Deep-Sea Drilling Project Leg-64, Gulf of California: Deep Sea

Drilling Project, Initial Reports, v. 64, p. 553–569.

KENDALL, C.G.ST.C., AND SKIPWITH, P., 1968, Recent algal mats of a Persian Gulf lagoon:

Journal of Sedimentary Petrology, v. 38, p. 1040–1058.

KIRKLAND, E.J., 1998, Advanced Computing in Electron Microscopy: New York, Plenum

Press, p. 29–59.

KITANO, Y., 1962, The behavior of various inorganic ions in the separation of calcium

carbonate from a bicarbonate solution: Chemical Society of Japan, Bulletin, v. 35, p.

1973–1980.

KNOLL, A.H., AND SWETT, K., 1990, Carbonate deposition during the late Proterozoic Era:

an example from Spitsbergen: American Journal of Science, v. 290A, p. 104–132.

KRAVCHINSKY, V.A., SKLYAROV, E.V., GLADKOCHUB, D.P., AND HARBERT, W.P., 2010,

Paleomagnetism of the Precambrian Eastern Sayan rocks: Implications for the

Ediacaran–Early Cambrian paleogeography of the Tuva–Mongolian composite terrane:

Tectonophysics, v. 486, p. 65–80.

KUNZMANN, M., HALVERSON, G.P., SOSSI, P.A., RAUB, T.D., PAYNE, J.L., AND KIRBY, J., 2013,

Zn isotope evidence for immediate resumption of primary productivity after Snowball

Earth: Geology, v. 41, p. 27–30.

LAND, L.S., 1998, Failure to precipitate dolomite at 258C from dilute solution despite 1000-

fold oversaturation after 32 years: Aquatic Geochemistry, v. 4, p. 361–368.

LAPORTE, L.F., 1967, Carbonate deposition near mean sea-level and resultant facies mosaic,

Manlius Formation (Lower Devonian) of New York State: American Association of

Petroleum Geologists, Bulletin, v. 51, p. 73–101.

LAST, F.M., LAST, W.M., AND HALDEN, N.M., 2010, Carbonate microbialites and

hardgrounds from Manito Lake, an alkaline, hypersaline lake in the northern Great

Plains of Canada: Sedimentary Geology, v. 225, p. 34–49.

LAST, F.M., LAST, W.M., AND HALDEN, N.M., 2012, Modern and late Holocene dolomite

formation: Manito Lake, Saskatchewan, Canada: Sedimentary Geology, v. 281, p. 222–

237.

LI, F.-B., TENG, F.-Z., CHEN, J.-T., HUANG, K.-J., WANG, S.-J., LANG, X.-G., MA, H.-R., PENG,

Y.-B., AND SHEN, B., 2016, Constraining ribbon rock dolomitization by Mg isotopes:

implications for the ‘‘dolomite problem’’: Chemical Geology, v. 455, p. 208–220.

LIBERMANN, O., 1967, Synthesis of dolomite: Nature, v. 213, p. 241–245.

LICHTENTHALER, H.K., STOBER, F., AND LANG, M., 1992, The nature of the different laser-

induced fluorescence signatures of plants: Advances in Remote Sensing, v. 1, p. 20–32.

LIPPMANN, F., 1973, The polymorphism calcite–aragonite, in Lippmann, L., ed.,

Sedimentary Carbonate Minerals: New York, Springer, p. 97–147.

LIU, C., WANG, Z., RAUB, T.D., MACDONALD, F.A., AND EVANS, D.A.D., 2014, Neoproterozoic

cap-dolostone deposition in stratified glacial meltwater plume: Earth and Planetary

Science Letters, v. 404, p. 22–32.

LOGAN, B.W., 1961, Cryptozoon and Associate Stromatolites from the Recent, Shark Bay,

Western Australia: The Journal of Geology, v. 69, p. 517–533.

LOKIER, S., AND STEUBER, T., 2008, Quantification of carbonate-ramp sedimentation and

progradation tates for the late Holocene Abu Dhabi shoreline: Journal of Sedimentary

Research, v. 78, p. 423–431.

MACHEL, H.G., 2004, Concepts and models of dolomitization: a critical reappraisal, in

Braithwaite, C.J.R., Rizzi, G., and Darke, G., eds., The Geometry and Petrogenesis of

Dolomite Hydrocarbon Reservoirs: Geological Society of London, Special Publication

235, p. 7–63.

MACHEL, H.G., AND MOUNTJOY, E.W., 1986, Chemistry and environments of dolomitization:

a reappraisal: Earth-Science Reviews, v. 23, p. 175–222.

MATTER, A., 1967, Tidal flat deposits in the Ordovician of western Maryland: Journal of

Sedimentary Petrology, v. 37, p. 601–609.

MAZZULLO, S.J., 2000, Organogenic dolomitization in peritidal to deep-sea sediments:

Journal of Sedimentary Research, v. 70, p. 10–23.

MAZZULLO, S.J., BISCHOFF, W.D., AND TEAL, C.S., 1995, Holocene shallow-subtidal

dolomitization by near-normal seawater, northern Belize: Geology, v. 23, p. 341–344.

MEYERS, S.R., AND SAGEMAN, B.B., 2007, Quantification of deep-time orbital forcing by

average spectral misfit: American Journal of Science, v. 307, p. 773–792.

MEYERS, S.R., SAGEMAN, B.B., AND ARTHUR, M.A., 2012, Obliquity forcing of organic

matter accumulation during Oceanic Anoxic Event 2: Paleoceanography, v. 27, PA3212.

MILORI, D.M.B.P., GALETI, H.V.A., MARTIN-NETO, L., DIECKOW, J., GONZALEZ-PEREZ, M.,

BAYER, C., AND SALTON, J., 2006, Organic matter study of whole soil samples using laser-

induced fluorescence spectroscopy: Soil Science Society of America, Journal, v. 70, p.

57–63.

MITCHELL, R.W., 1985, Comparative sedimentology of shelf carbonates of the Middle

Ordovician St. Paul Group, central Appalachians: Sedimentary Geology, v. 43, p. 1–41.

MORRIS, R.V, RUFF, S.W., GELLERT, R., MING, D.W., ARVIDSON, R.E., CLARK, B.C., GOLDEN,

D.C., SIEBACH, K., KLINGELHOFER, G., SCHRODER, C., FLEISCHER, I., YEN, A.S., AND

SQUYRES, S.W., 2010, Identification of carbonate-rich outcrops on Mars by the Spirit

Rover: Science, v. 329, p. 421–424.

MORRIS, R.V., BLAKE, D.F., BISH, D., MING, D.W., AGRESTI, D.G., TRIEIMAN, A.H., STEELE,

A., AMUNDSEN, H.E.F., AND TEAM, A., 2011, A terrestrial analogue from Spitsbergen

(Svalbard, Norway) for the Comanche carbonate at Gusev Crater, Mars: Lunar and

Planetary Science, v. 42, p. 1699–1700.

MORSE, J.W., WANG, Q., AND TSIO, M.Y., 1997, Influences of temperature and Mg: Ca ratio

on CaCO3 precipitates from seawater: Geology, v. 25, p. 85–87.

MUCCI, A., AND MORSE, J.W., 1983, The incorporation of Mg2þ and Sr2þ into calcite

overgrowths: influences of growth rate and solution composition: Geochimica et

Cosmochimica Acta, v. 47, p. 217–233.

NAVROTSKY, A., DOOLEY, D., REEDER, R., AND BRADY, P., 1999, Calorimetric studies of the

energetics of order–disorder in the system Mg1–xFexCa(CO3)2: American Mineralogist,

v. 84, p. 1622–1626.

NELLIST, P.D., 2007, Scanning transmission electron microscopy, in Hawkes, P.W., and

Spence, J.C.H., eds., Science of microscopy: New York, Springer, p. 65–132.

NEUMANN, A.C., GEBELEIN, C.D., AND SCOFFIN, T.P., 1969, Composition, Structure, and

Erodability of Subtidal Mats, Abaco, Bahamas [Abstract]: American Association of

Petroleum Geologists, Bulletin, v. 53, p. 734.

OOMORI, T., AND KITANO, Y., 1987, Synthesis of protodolomite from sea water containing

dioxane: Geochemical Journal, v. 21, p. 59–65.

PAULL, C.K., NEUMANN, A.C., BEBOUT, B., ZABIELSKI, V., AND SHOWERS, W., 1992, Growth

rate and stable isotopic character of modern stromatolites from San Salvador, Bahamas:

Palaeogeography, Palaeoclimatology, Palaeoecology, v. 95, p. 335–344.

PETRASH, D.A., GUENELI, N., BROCKS, J.J., MENDEZ-DOT, J.A., GONZALEZ-ARISMENDI, G.,

POULTON, S.W., AND KONHAUSER, K.O., 2016, Black shale deposition and early diagenetic

dolomite cementation during Oceanic Anoxic Event 1: the mid-Cretaceous Maracaibo

Platform, northwestern South America: American Journal of Science, v. 316, p. 669–

711.

PETRASH, D.A., BIALIK, O.M., BONTOGNALI, T.R.R., VASCONCELOS, C., ROBERTS, J.A.,

MCKENZIE, J.A., AND KONHAUSER, K.O., 2017, Microbially catalyzed dolomite formation:

from near-surface to burial: Earth-Science Reviews, v. 171, p. 558–582.

PISCIOTTO, K.A., SEA, D., AND JOLLA, L., 1981, Isotopic survey of diagenetic carbonates:

Deep Sea Drilling Project Leg 63: Deep Sea Drilling Project, Initial Reports, v. 63, p.

595–609.

RAZ, S., WEINER, S., AND ADDADI, L., 2000, Formation of high-magnesian calcites via an

amorphous precursor phase: possible biological implications: Advanced Materials, v. 12,

p. 38–42.

READ, J.F., 1980, Carbonate ramp-to-basin transitions and foreland basin evolution, Middle

Ordovician, Virginia Appalachians: American Association of Petroleum Geologists,

Bulletin, v. 64, p. 1575–1612.

REDDY, M.M., 1977, Crystallization of calcium carbonate in the presence of trace

concentrations of phosphorus-containing anions: I. Inhibition by phosphate and

glycerophosphate ions at pH 8.8 and 25 C: Journal of Crystal Growth, v. 41, p. 287–295.

REDDY, M.M., AND NANCOLLAS, G.H., 1976, The crystallization of calcium carbonate: IV.

The effect of magnesium, strontium and sulfate ions: Journal of Crystal Growth, v. 35, p.

33–38.

REEDER, R.J., AND SHEPPARD, C.E., 1984, Variation of lattice parameters in some

sedimentary dolomites: American Mineralogist, v. 69, p. 520–527.

REEDER, R.J., AND WENK, H.R., 1983, Structure refinements of some thermally disordered

dolomites: American Mineralogist, v. 68, p. 769–776.

ROBERTS, J.A., KENWARD, P.A., FOWLE, D.A., GOLDSTEIN, R.H., GONZALEZ, L.A., AND MOORE,

D.S., 2013, Surface chemistry allows for abiotic precipitation of dolomite at low

temperature: National Academy of Sciences [USA], Proceedings, v. 110, p. 14,540–

14,545.

Y. FANG AND H. XU694 J S R

RUSHDI, A.I., PYTKOWICZ, R.M., SUESS, E., AND CHEN, C.T., 1992, The effects of magnesium-

to-calcium ratios in artificial seawater, at different ionic products, upon the induction

time, and the mineralogy of calcium carbonate: a laboratory study: Geologische

Rundschau, v. 81, p. 571–578.

SANDER, B., 1936, Contributions to the study of Depositional Fabrics: Rhythmically

Deposited Triassic Limestones and Dolomites (translated into English by Knopf, E.B.,

1951): American Association of Petroleum Geologists, 207 p.

SANDO, W.J., 1957, Beekmantown Group (Lower Ordovician) of Maryland: Geological

Society of America, Memoir 68, 161 p.

SHEN, Z., KONISHI, H., BROWN, P.E., AND XU, H., 2013, STEM investigation of exsolution

lamellae and ‘‘c’’ reflections in Ca-rich dolomite from the Platteville Formation, western

Wisconsin: American Mineralogist, v. 98, p. 760–766.

SHEN, Z., KONISHI, H., SZLUFARSKA, I., BROWN, P.E., AND XU, H., 2014, Z-contrast imaging

and ab initio study on ‘‘D’’ superstructure in sedimentary dolomite: American

Mineralogist, v. 99, p. 1413–1419.

SHEN, Z., SZLUFARSKA, I., BROWN, P.E., AND XU, H., 2015, Investigation of the role of

polysaccharide in the dolomite growth at low temperature by using atomistic simulations:

Langmuir, v. 31, p. 10,435–10,442.

SHIELDS, G.A., 2005, Neoproterozoic cap carbonates: a critical appraisal of existing models

and the plumeworld hypothesis: Terra Nova, v. 17, p. 299–310.

SHINN, E.A., LLOYD, R.M., AND GINSBURG, R.N., 1969, Anatomy of a modern carbonate

tidal-flat, Andros Island, Bahamas: Journal of Sedimentary Petrology, v. 39, p. 1202–

1228.

SIBLEY, D.F., AND GREGG J.M., 1987, Classification of dolomite rock textures: Journal of

Sedimentary Petrology, v. 57, p. 967–975.

SONE, H., AND ZOBACK, M.D., 2013, Mechanical properties of shale-gas reservoir rocks:

Part 2: ductile creep, brittle strength, and their relation to the elastic modulus:

Geophysics, v. 78, p. 393–402.

SONNENBERG, S.A., LEFEVER, J.A., AND HILL, R.J., 2011, Fracturing in the Bakken petroleum

system, Williston Basin, in Robinson, J.W., LeFever, J.A., and Gaswirth, S.B., eds., The

Bakken–Three Forks Petroleum System in the Williston Basin: Rocky Mountain

Association of Geologists, p. 393–417.

STEPHENSON, A.E., DEYOREO, J.J., WU, L., WU, K.J., HOYER, J., AND DOVE, P.M., 2008,

Peptides enhance magnesium signature in calcite: insights into origins of vital effects:

Science, v. 322, p. 724–727.

TIANO, P., ACCOLLA, P., AND TOMASELLI, L., 1995, Phototrophic biodeteriogens on lithoid

surfaces: an ecological study: Microbial ecology, v. 29, p. 299–309.

TUCKER, M.E., AND WRIGHT, V.P., 1990, Carbonate sedimentology: Malden, MA, Blackwell

Publishing, 482 p.

VAN LITH, Y., VASCONCELOS, C., WARTHMANN, R., MARTINS, J.C.F., AND MCKENZIE, J.A.,

2002, Bacterial sulfate reduction and salinity: two controls on dolomite precipitation in

Lagoa Vermelha and Brejo do Espinho (Brazil): Hydrobiologia, v. 485, p. 35–49.

VAN LITH, Y., WARTHMANN, R., VASCONCELOS, C., AND MCKENZIE, J.A., 2003, Microbial

fossilization in carbonate sediments: a result of the bacterial surface involvement in

dolomite precipitation: Sedimentology, v. 50, p. 237–245.

VAN TUYL, F.M., 1914, The origin of dolomite: Iowa Geological Survey, Annual Report, v.

XXV, p. 251–421.

VASCONCELOS, C., AND MCKENZIE, J.A., 1997, Microbial mediation of modern dolomite

precipitation and diagenesis under anoxic conditions (Lagoa Vermelha, Rio de Janeiro,

Brazil): Journal of Sedimentary Research, v. 67, p. 378–390.

WANLESS, H.R., 1979, Limestone response to stress: pressure solution and dolomitization:

Journal of Sedimentary Petrology, v. 49, p. 437–462.

WARREN, J., 2000, Dolomite: occurrence, evolution and economically important

associations: Earth-Science Reviews, v. 52, p. 1–81.

WILLIAMS, G.E., GOSTIN, V.A., MCKIRDY, D.M., PREISS, W. V., AND SCHMIDT, P.W., 2011, The

Elatina glaciation (late Cryogenian), South Australia, in Arnaud, E., Halverson, G.P., and

Shields-Zhou, G., eds., Geological Society of London, Memoir 36, p. 713–721.

WRIGHT, D.T., AND WACEY, D., 2005, Precipitation of dolomite using sulphate-reducing

bacteria from the Coorong Region, South Australia: significance and implications:

Sedimentology, v. 52, p. 987–1008.

XU, J., YAN, C., ZHANG, F., KONISHI, H., XU, H., AND TENG, H.H., 2013, Testing the cation-

hydration effect on the crystallization of Ca–Mg–CO3 systems: National Academy of

Sciences [USA], Proceedings, v. 110, p. 17750–17755.

XU, H., SHEN, Z., KONISHI, H., AND LUO, G., 2014, Crystal structure of Guinier-Preston

zones in orthopyroxene: Z-contrast imaging and ab inito study: American Mineralogist,

v. 99, p. 2043–2048.

ZENGLER, D.H., DUNHAM, J.D., AND ETHINGTON, R.L., eds., 1980, Concepts and Models of

Dolomitization: SEPM, Special Publication 28, 320 p.

ZENTMYER, R.A., PUFAHL, P.K., JAMES, N.P., AND HIATT, E.E., 2011, Dolomitization on an

evaporitic Paleoproterozoic ramp: widespread synsedimentary dolomite in the Denault

Formation, Labrador Trough, Canada: Sedimentary Geology, v. 238, p. 116–131.

ZHANG, F., XU, H., KONISHI, H., AND RODEN, E.E., 2010, A relationship between d104 value

and composition in the calcite-disordered dolomite solid-solution series: American

Mineralogist, v. 95, p. 1650–1656.

ZHANG, F., XU, H., KONISHI, H., KEMP, J.M., RODEN, E.E., AND SHEN, Z., 2012a, Dissolved

sulfide-catalyzed precipitation of disordered dolomite: implications for the formation

mechanism of sedimentary dolomite: Geochimica et Cosmochimica Acta, v. 97, p. 148–

165.

ZHANG, F., XU, H., KONISHI, H., SHELOBOLINA, E.S., AND RODEN, E.E., 2012b,

Polysaccharide-catalyzed nucleaction and growth of disordered dolomite: a potential

precursor of sedimentary dolomite: American Mineralogist, v. 97, p. 556–567.

ZHANG, F., YAN, C., TENG, H.H., RODEN, E.E., AND XU, H., 2013, In situ AFM observations

of Ca-Mg carbonate crystallization catalyzed by dissolved sulfide: implications for

sedimentary dolomite formation: Geochimica et Cosmochimica Acta, v. 105, p. 44–55.

ZHANG, F., XU, H., SHELOBOLINA, E.S., KONISHI, H., CONVERSE, B., SHEN, Z., AND RODEN,

E.E., 2015, The catalytic effect of bound extracellular polymeric substances excreted by

anaerobic microorganisms on Ca-Mg carbonate precipitation: implications for the

‘‘dolomite problem’’: American Mineralogist, v. 100, p. 483–494.

Received 7 November 2017; accepted 20 March 2018.

MICROBIAL IMPACT ON SEDIMENTARY DOLOMITE FORMATION FROM A LAMINATED CARBONATEJ S R 695