Received Date: 24-Feb-2015

Revised Date: 25-Feb-2015

Accepted Date: 16-Mar-2015

Article Type: Original Manuscript

The Geochemistry of Carbonate Diagenesis: The Past, Present and Future

Peter K. Swart,

Department of Marine Geosciences,

Rosenstiel School of Marine and Atmospheric Sciences,

University of Miami, Miami Fl 33149.

pswart@rsmas.miami.edu

Short Title – Carbonate diagenesis

Keywords – Carbonate, diagenesis, dolomite, geochemistry, isotopes

ABSTRACT

Stable carbon and oxygen isotopes (18O and 13C values) and trace elements have been

applied to the study of diagenesis of carbonate rocks for over 50 years. As valuable as these

insights have been, many problems regarding the interpretation of geochemical signals within

mature rocks remain. For example, while the 18O values of carbonate rocks are dependent

both upon the temperature and the 18O value of the fluid and additional information,

including trace element composition aids in interpreting such signals, direct evidence of either

the temperature or the composition of the fluids is required. Such information can be obtained
This is an Accepted Article that has been peer-reviewed and approved for publication in the
Sedimentology, but has yet to undergo copy-editing and proof correction. Please cite this article
as an “Accepted Article”; doi: 10.1111/sed.12205

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by analyzing the 18O value of any fluid inclusions or by measuring the temperature using a

method such as the ‘clumped’ isotope technique. Such data speak directly to a large number of

problems in interpreting the oxygen isotope record including the well-known tendency for 18O

values of carbonate rocks to decrease with increasing age. Unlike the 18O, 13C values of

carbonates are considered to be less influenced by diagenesis and more a reflection of primary

changes in the global carbon cycle through time. However, many studies have not sufficiently

emphasized the effects of diagenesis and other post-depositional influences on the eventual

carbon isotopic composition of the rock with the classic paradigm that the present is the key to

the past being frequently ignored. Finally, many additional proxies are poised to contribute to

the interpretation of carbonate diagenesis. Although the study of carbonate diagenesis is at an

exciting point with an explosion of new proxies and methods, care should be taken to

understand both old and new proxies before applying them to the ancient record.

INTRODUCTION

Although geochemical changes during the diagenesis of carbonate sediments and rocks have

been recognised for a considerable period of time, the development and refinement of the

techniques for the analysis of stable isotopes of oxygen (18O and 16O) and carbon (13C and 12C)

(18O and 13C values) in carbonates (Epstein et al., 1951; Epstein et al., 1953; McCrea, 1950)

have provided insights into the use of isotopes as geochemical tracers in the interpretation of

alteration (diagenesis) of carbonate sediments. Early papers (Gross, 1964; Gross & Tracey,

1966; Hudson, 1977; Land, 1967; Matthews, 1968) provided the foundation for subsequent

work which refined and explored new interpretations and applications. In addition, the

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increased precision and ease of the elemental analyses of carbonates has led to the use of a

combination of trace elements (Banner, 1995; Brand & Veizer, 1980; Brand & Veizer, 1981;

Kinsman, 1973) and stable C and O isotopes to unravel diagenetic histories in carbonate rocks.

There have been several important and valuable texts dealing with this topic (Bathurst, 1971;

Berner, 1980; Milliman, 1974; Moore, 1989; Morse & Mackenzie, 1990) as well as numerous

review papers (Budd, 1997; Drever, 1982; Emerson & Hedges, 2003; Holland, 2007; Lerman &

Clauer, 2007; Martin & Sayles, 2003; Morse, 2003; Veizer & Mackenzie, 2003) and others.

Background on Diagenetic Geochemical Indices

Carbonates are considered to include any mineral with a structural CO32- group. Although many

metals form carbonate minerals, the most common carbonates are those formed by Ca, Mg,

and a combination of these two elements. The basic distribution of carbon and oxygen

isotopes in nature is well covered in numerous publications and text books (Friedman & O'Neil,

1977; Hoefs, 1980; Sharp, 2007; Zeebe & Wolf-Gladrow, 2001) and will not be considered here

in detail. However, a brief introduction has been included in order to guide the reader to

appropriate references and to provide a historical context.

Oxygen

The 18O value of carbonate minerals is controlled by: (i) the temperature of formation (Epstein

et al., 1951; Friedman and O'Neil, 1977; O'Neil et al., 1969; Tremaine et al., 2011; Urey, 1947);

(ii) the 18O value (18Ow) of the precipitating fluid (Epstein et al., 1953; Epstein & Mayeda,

1953; Urey, 1947); (iii) mineralogy (Emrich et al., 1970; Tarutani et al., 1969); (iv) the pH of the

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solution (Deuser & Degens, 1967; Zeebe & Wolf-Gladrow, 2001); and (v) any kinetic effects

manifested during the precipitation procedure (McConnaughey, 2003).

Temperature

Since the paper of Urey (1947), the focus of the 18O technique has been mainly as a

temperature proxy. The 18O value of a carbonate mineral is inversely correlated to

temperature and for those carbonates forming at sedimentary temperatures the relationship is

approximately 4oC for every change of 1‰ in the 18O value.

Fluids

The 18Ow is controlled by the origin of rainfall, the amount of evaporation (Craig & Gordon,

1965; Craig et al., 1963; Gonfiantini, 1986), mixing with seawater and mineral–water reactions

(Lawrence, 1989) . Typically meteoric waters have 18Ow values less than 0‰ and these

become more negative towards polar regions, with increasing altitude, distance from the

evaporation source and/or decreasing temperature (Rozanski et al., 1993; Fig. 1).

Compared to meteoric fluids marine waters have more positive 18Ow values and, although

during evaporation water bodies becomes more enriched in 18O, in practice there is a limit to

the extent to which 18Ow values can increase because the 18O and 16O in the water vapour in

the atmosphere isotopically exchanges with the evaporating water pool (Craig & Gordon, 1965).

In addition, the 18Ow value can become altered by changes in the thermodynamic activity of

water, as a consequence of the hydration of ions in solution, and from isotope fractionation

during the crystallization of salts (Gonfiantini, 1986). As a result, during the final stages of

evaporation the 18Ow value may reach a maximum value and then become more negative as

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evaporation proceeds. The ultimate 18Ow value of an evaporating fluid is therefore tied to the

relative humidity of the atmosphere, the temperature of evaporation, the 18O of the water

vapour and the initial chemical composition of the fluid.

Mineralogy

Inorganically formed aragonite (Grossman, 1984) and dolomite (Fritz & Smith, 1970; Matthews

& Katz, 1977; Northrop & Clayton, 1966; O'Neil et al., 1969; O'Neil & Epstein, 1966 ; Sheppard &

Schwarcz, 1970; Vasconcelos et al., 2005) have different equations compared to low Mg calcite

that relate temperature to the 18O value of the mineral. The 18O of aragonite, for example, is

typically about 1‰ more positive than that of low Mg calcite (LMC) formed at the same

temperature, while the 18O value of high Mg calcite (HMC) increases by about 0.06‰ for

every mol% MgCO3 (Tarutani et al., 1969). Interestingly, the 0.06‰ difference corresponds to

the theoretical and expected difference of 3‰ between the 18O values of calcite and dolomite

at 25oC (Land, 1980; Land, 1983; see later discussion on dolomite).

Kinetics and pH

The distribution of the inorganic oxygen bearing carbon species in solution is controlled by pH.

At high pH the speciation is dominated by CO32- changing to HCO3- and H2CO3 as the pH is

lowered. Because there are well-known isotopic fractionations for both C and O between these

species (Beck et al., 2005; Emrich et al., 1970; Zeebe & Wolf-Gladrow, 2001), the 18O value of

the sum of these species changes as a function of pH assuming that pH is the only factor driving

the speciation (Zeebe, 2005a; Zeebe, 2007; Zeebe & Wolf-Gladrow, 2001). At low pH the major

species, H2CO3, is about 41 ‰ more positive than the oxygen in the water at 25oC (Bottinga &

Craig, 1969). As pH increases and HCO3- and CO32- become the dominant species, the 18O

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value of the sum of the carbonate species becomes more negative. Assuming that the 18O

value of carbonate precipitated reflects the 18O value of the sum of the carbonate species and

that equilibrium is not maintained between the various species (Zeebe, 2005a), then at low pH

it would be expected that carbonates would have more positive 18O values than those

precipitated at high pH (Zeebe & Wolf-Gladrow, 2001; Fig. 2).

Kinetic controls are considered together with pH because the rate of precipitation is often

related to the amount of CO32- in a solution which in turn is controlled by the pH. Generally

faster rates of precipitation will occur at higher pH and favour the incorporation of the light

isotopes, 16O and 12C relative to 18O and 13C. Kinetic and pH controls are particularly important

in the precipitation of biogenic carbonates, but also play a role during the precipitation of some

inorganic carbonates (i.e. speleothems; Affek et al., 2008b; Tremaine et al., 2011).

Carbon

The 13C values in carbonates are controlled by the: (i) 13C values of ambient dissolved

inorganic carbonate (DIC; Mook, 1968); (ii) pH of precipitation (Zeebe & Wolf-Gladrow, 2001);

(iii) rate of precipitation (kinetics; McConnaughey, 2003); (iv) mineralogy (Emrich et al., 1970;

Rubinson & Clayton, 1969); and (v) temperature (Deines et al., 1974; Emrich et al., 1970).

Dissolved Inorganic Carbon

The major control on the 13C values of marine carbonates is exerted by the 13C value of

oceanic DIC; this in turn is determined by the fixation of CO2 during photosynthesis (Craig,

1953; Park & Epstein, 1961), the weathering of carbonate rocks (Berner et al., 1983) and the

oxidation of organic material (OM) within the marine environment (Weber & Woodhead,

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1971). The extent to which OM is depleted in 13C depends upon the photosynthetic pathway

utilized by the plant (C3, C4 or CAM.) Most photosynthetic terrestrial vegetation utilizes the C3

pathway and has a 13C value between -25‰ and -30‰, while marine vascular plants utilizing

the same pathway have significantly more positive values (-8 to -12‰; Deines, 1980). The

differences in the 13C values of aqueous and terrestrial vascular plants reflects the fact that the

majority of inorganic carbon in the aqueous realm is usually present as HCO3-, and is ca 8‰

enriched in 13C compared to atmospheric CO2 (Vogel et al., 1970). In contrast to marine

vascular plants, marine algae have much more negative 13C values (-20 to -24‰; Degens et al.,

1968; Sackett et al., 1965). Variations in the 13C values of DIC are relatively small in the open

ocean, but become increasingly important in coastal marine environments, such as coral reefs,

isolated embayments and shallow carbonate shelves, where the residence time of the water is

longer (Weber & Woodhead, 1971). The 13C values of DIC in these environments can

therefore be influenced by the addition of isotopically negative CO2 from respiration as well as

fractionation and enrichment in the 13C of the inorganic pool by photosynthesis.

Kinetics and pH

The pH and kinetic control of carbon in water operates in a similar manner to that of oxygen

except that the pool of carbon is much smaller. As a result, the range of 13C values over which

the different species vary is restricted compared to that of 18O. Consequently, between pH

values of ca 4 to 10 there is a strong covariance between 13C and 18O values in carbonates

affected by changes in pH (Zeebe & Wolf-Gladrow, 2001). Above pH values of ca 10 the 13C

value of the carbon species no longer varies, while the 18O value of the sum of the carbonate

species can still change to more negative values (Fig. 2).

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Mineralogy

The 13C value of aragonite is typically 1 to 2‰ more positive than co-occurring LMC and HMC

(Emrich et al., 1970; Romanek et al., 1992; Rubinson & Clayton, 1969). Dolomites may be ca

1‰ more positive than LMC formed under similar conditions (Sheppard & Schwarcz, 1970).

Temperature

In contrast to 18O, the 13C value of carbonates is not thought to be directly influenced by

temperature. However, it is clear that the relationships between the various inorganic bearing

species are temperature-dependent (Deines et al., 1974; Emrich et al., 1970). Therefore, in

spite of experimental data which suggests that the 13C values of carbonates is independent of

temperature (Romanek et al., 1992; Rubinson & Clayton, 1969), theoretically the 13C value of

carbonate should change. Part of the problem may be the limited number of studies measuring

the fractionation factors between various species and the inconsistent values reported. For

example, in the study of Deines et al. (1974), the fractionation between calcite and HCO3-

decreased with decreasing temperature (contrary to the expected pattern), while the data

presented by Emrich et al. (1970) were based on carbonates of mixed aragonite and calcite

mineralogy. Regardless of whether the 13C values of inorganic precipitates are temperature

dependent, because processes of respiration and photosynthesis are related to temperature,

temperature will play an important role in governing the 13C value of DIC and therefore

ultimately the 13C value of the precipitated carbonate.

Clumped Isotopes

The major problem with using 18O values in minerals such as carbonates is that the values

depend both upon temperature and the 18Ow value (Fig. 3). This dilemma is potentially

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solvable using the 47, which allows formation or alteration temperatures to be measured

independently of the 18Ow. This method has already been successfully applied to a wide range

of previously used palaeotemperature proxies (Affek et al., 2008a; Affek et al., 2008b; Eagle et

al., 2011; Ghosh et al., 2007a; Ghosh et al., 2007b; Huntington et al., 2011; Peters et al., 2012;

Snell et al., 2007; Thiagarajan et al., 2011; Tripati et al., 2010; Wacker et al., 2014). Although

the application of the clumped isotope technique to diagenetic studies is still in the early

stages, several papers (Bristow et al., 2011; Budd et al., 2013; Dennis & Schrag, 2010;

Fernandez et al., 2014; Ferry et al., 2011; Loyd et al., 2012; Loyd et al., 2013; Sena et al., 2014;

Van De Velde et al., 2013) have suggested that this could be a promising approach. However,

there have been no studies examining the behaviour of clumped isotopes in well-constrained

diagenetic settings. The original calibration (Ghosh et al., 2006), applied to a range of different

carbonates, has been augmented by a theoretical calibration (Guo et al., 2009) and other

empirical calibrations (Dennis & Schrag, 2010). The slight differences in slopes and intercepts

between these are sufficient to cause significant error in the calculated temperature. The

promise of the clumped isotope method is that once the temperature has been calculated, the

18Ow value of the formation or diagenetic fluid can be estimated using known relationships

between the fluid and mineral. Although the uncertainty in the various equations relating 47

to temperature is still significant, reasonable estimates can be made using the equations

published for LMC.

Solid-state Resetting

The clumped isotope method also allows investigation of the phenomenon of isotopic resetting

in carbonates at temperatures higher or lower than those of formation or diagenetic alteration.

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For example, does a calcite or dolomite mineral formed at sedimentary temperatures, change

its ratio of 13C–18O bonds at higher temperature and, if so, at what temperature and at what

rate? Alternatively, do minerals formed at high temperatures continually reorganize their 13C–
18
O bonds until a ratio is locked in at lower temperatures? Such a phenomenon would be

important because it would constrain the burial history of a carbonate and perhaps the original

depositional temperatures could be back calculated assuming that there was information on

the rate of change at specific temperatures and the burial temperature. If solid diffusion was

responsible for these changes, then the 18Ow calculated using such temperatures would in fact

be erroneous unless the temperatures were corrected for the burial history. Studies

investigating this phenomenon indicate that such resetting does in fact take place at

geologically significant rates, particularly at temperatures higher than 300oC (Henkes et al.,

2014; Passey & Henkes, 2012).

Carbon and Oxygen Isotopic Variations in Modern Carbonates

Dissolved Inorganic Carbon in the Oceans

During periods of enhanced preservation and burial of organic matter (OM) the 13C value of

the DIC increases. During the Early Carboniferous for example, when abundant coal deposits

were formed, the 13C value of the DIC in the oceans, as recorded by carbonate minerals,

increased by 3 to 4‰ (Saltzman et al., 2004; Veizer et al., 1999). Conversely, during times of

enhanced oxidation of OM, the 13C value of the DIC decreases. Over the past 50 Myr, the 13C

values of pelagic carbonates have decreased by 2 to 3‰ (Shackleton, 1985; Tipple et al., 2010)

reflecting the transfer of organic carbon into the ocean-atmosphere system. Enhanced

oxidation can also occur during periods of low sea level, when the continental shelves are

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exposed. Lower sea level can also destabilize gas hydrates, which form in sediments on the

continental margin, thereby releasing large volumes of CH4 with low 13C values into the

atmosphere. Such CH4 ultimately oxidizes to CO2, causing the atmosphere and the oceanic DIC

to return to more negative values. Similar changes can also be associated with extreme events

such as rapid warming, extra-terrestrial impacts, or man-made events such as the burning of

fossil fuel.

Marine Cements

Inorganically precipitated carbonates which form in equilibrium with their environment possess

18O and 13C values which agree with those calculated from theoretically and/or empirically

derived equations. Because different types of carbonate (HMC, low-Mg LMC, dolomite and

aragonite) have distinctive equations which describe equilibrium, these differences will be

evident in the 18O and 13C values. Differences in the 18O and 13C values of carbonate with

varying mineralogy is confirmed by analyses of cements from modern environments, such as

the aragonites which form in the cavities of steep marine slopes in the Bahamas (Gonzalez &

Lohmann, 1985; Grammer et al., 1993) or HMC cements which are found within coral reefs in

the Pacific (Aissaoui, 1988; Gonzalez & Lohmann, 1985). The 13C values of these cements are

slightly more negative than those of inorganic aragonite precipitated on margins of the Great

Bahama Bank (GBB), perhaps as a result of microbial processes occurring within the interstices

of the reef framework (Camoin & Seard, 2012; Fig. 5).

Biogenic Carbonates

Biogenically produced carbonates can either be precipitated in isotopic equilibrium with their

environments or can exhibit so-called ‘vital effects’ in which the 18O and 13C values are

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shifted from equilibrium, typically towards more negative values. Variations in the ranges of

18O and 13C values in carbonate secreting organisms arise as a result of differences in the

mechanism of calcification, mineralogy, presence of photosymbionts and the rate of

calcification (Adkins et al., 2003; McConnaughey, 1989; McConnaughey, 2003). The 18O and

13C values of such skeletal material have been reported in numerous publications and can vary

widely (Fig. 6). As an example, aragonite from photosymbiont-bearing, shallow water coral

skeletons typically have 18O values between -5‰ and -4‰ and 13C values between -1‰ and

+1‰ (Swart, 1983) . Such 18O values are too negative to be indicative of equilibrium in the

environments in which coral reefs are normally found. The 13C values of the skeletons are also

too variable when compared to the relatively constant 13C values of the ambient DIC. In

contrast to shallow-water corals, deep-water varieties, which lack photosymbionts, have 18O

values between -8‰ and -2‰ and 13C values between -8‰ and +1‰ (Adkins et al., 2003;

Emiliani et al., 1978; Land et al., 1977; Fig. 6). These wide ranges in the deep-water corals

cannot be caused by changes in the 18O and 13C values of the water or the temperature of the

environment and therefore they must be produced by changes in the calcification process, such

as calcification rate (McConnaughey, 2003), pH (Adkins et al., 2003), or both. In contrast, the

photosymbionts in the shallow-water corals modify the calcification environment, limiting wide

variations in pH during calcification and therefore have more restrictive 18O and 13C ranges

compared to non-photosymbiont bearing corals. Insight into the behaviour of the 18O and

13C values of the non-symbiont bearing corals can be obtained by comparing data with the

theoretical behaviour of 18O and 13C values as a function of pH shown in Fig. 2.

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Other calcareous organisms such as red and green algae show distinctive 18O and 13C ranges

relating to the calcification processes (Figs 5 and 6). Green algae tend to have more positive

13C values, although different portions of the algae can have different compositions. For

example the head of algae such as Penicillus sp. and Rhipocephalus sp. tend to have more

positive 18O and 13C values compared to the stalk. In Halimeda sp. the 13C values of the

newly formed algal segments can be up to 6‰ heavier than the basal portion (Wefer, 1981). A

similar enrichment was observed in Penicillus sp. by these authors. Although the authors

offered no explanation for the observed pattern in 13C values, it is likely that it results from

secondary calcification within the algal segments as they age. In contrast calcareous red algae

are much more depleted in 13C than green algae and can be considered to form out of isotopic

equilibrium (Keith & Weber, 1965).

Organic Material within Modern Carbonates

Most modern carbonates, whether they are considered non-biogenic (ooids and peloids) or

produced by calcareous organisms, contained variable amounts of OM (Ingalls et al., 2003;

Muscatine et al., 2005; Wainwright, 1964). Such OM can be produced by the organism itself as

a result of the calcification process, or be secondary as a result of the action of algae, fungi,

bacteria, sponges and other boring organisms. Such secondary OM contributes significantly to

the total organic content of many skeletons with, for example, boring organisms observed in

intimate contact with recently formed coral skeletons (Golubic, 1969; Lukas, 1974). In fact it is

probable that the OM originating from secondary contributors is greater than that supplied by

the host organism. The pervasiveness of such material can be observed readily through the use

of thin sections and scanning electron microscopy (SEM) in live and recently dead skeletons

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(Fig. 7). Within the relatively closed environments of the interiors of skeletons, the OM is

subject to degradation causing lower levels of oxygen and possible sulphate reduction. The

importance of such material is that its 15N and 13C value has been used as a proxy for a

variety of different environmental parameters (Marion et al., 2006; Wang et al., 2014). Non-

skeletal materials such as ooids and peloids contain higher amounts of OM mainly derived from

endolithic organisms. Processes such as sulphate reduction and denitrification have been

identified in these particles using genetic markers (Diaz et al., 2014) and confirmed in the

geochemical signatures of pore water leachates derived from these particles (Diaz et al., 2013).

Carbonate Platforms

As a result of varying carbonate contributions derived or produced by a number of different

organisms and precipitation mechanisms (either organically or inorganically produced),

different carbonate environments tend to have characteristic 18O and 13C values. For

example, within a reef, the 18O and 13C values of the sediments change between different

sub-environments (i.e. reef flat, lagoon, reef crest and fore reef; Weber & Woodhead, 1969)

reflecting the varying contributions of different organisms. However, environments dominated

by non-skeletal grains (peloids, ooids, etc.), such as the Bahamas, tend to have more positive

18O and 13C values (18O = ca 0‰ and 13C = ca +3 to +5‰) and appear to be in isotopic

equilibrium (Lowenstam & Epstein, 1957; Shinn et al., 1989; Swart et al., 2009) with ambient

waters (Gischler et al., 2009; Weber & Woodhead, 1969; Fig. 5). For comparison, other skeletal

and non-skeletal components are shown in Fig. 6. The nature of the basin (closed or open) and

its connection to the open marine environment also affects the 18O and 13C values of the

waters and hence the values of carbonates precipitated from those waters (Patterson & Walter,

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1994). Semi-enclosed basins such as Florida Bay tend to have relatively positive 18Ow values

(Lloyd, 1964; Swart & Price, 2002) as a result of evaporation. The 13C values, however, will be

lowered following input of waters containing organic carbon or the products of the oxidation of

organic carbon from adjacent terrestrial areas, which are depleted in 13C (Halley & Roulier,

1999). A similar scenario prevails in ramp settings, both in modern and ancient carbonates

(Holmden et al., 1998; Katz et al., 2007).

Freshwater Carbonates

In addition to biogenic carbonates found in freshwaters, the classification scheme of freshwater

carbonates includes travertines (generally used for carbonates deposited in higher temperature

regimes related to hydrothermal activity), tufas (lower temperature deposits formed in lakes,

springs and waterfalls) and speleothems (precipitates formed in cave systems; Capezzuoli et al.,

2014; Ford & Pedley, 1996; Fouke, 2011; Sanders & Friedman, 1967). In contrast to marine

settings, freshwater carbonates, whether organic or inorganic in nature, generally have more

negative 18O and 13C values. The 18O values tend to be related to water temperature and

the values of the local meteoric fluids, with higher latitudes and altitudes having more negative

18Ow values that closely follow those of global precipitation patterns. The only exception to

the more negative patterns in 18Ow values would be terminal lakes such as Pyramid Lake

(Benson et al., 2013) and marsh areas such as the Everglades (Meyers et al., 1993; Price &

Swart, 2006) where the principal water loss is through evaporation (Gonfiantini, 1986; Leng &

Marshall, 2004; Yuan et al., 2011). Evaporation would produce freshwater carbonates with

more positive 18Ow values than expected considering the low salinity of the water bodies.

Such reverse patterns in the 18Ow values are observed in the transition between freshwater

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marshes such as the Everglades and adjacent marine areas (Swart & Price, 2002), and are

reflected in elevated 18O values in carbonates found in the region (Halley & Roulier, 1999;

Lloyd, 1964). This is opposite to the expected trend if a major river such as the Mississippi with

negative 18Ow values flowed into a marine estuary with more positive 18Ow values. In the

ancient a similar interpretation to that used for the Everglades has been placed on the 18O

values in Jurassic molluscs (Hendry & Kalin, 1997) where specimens with the most negative

18O values were found furthest from shore. The 13C values of lake carbonates are more

negative than marine carbonates, a phenomenon linked to the quantity of OM relative to the

size of the water body, its 13C value and subsequent oxidation. In addition, many tufas are

associated with microbial activity (Chafetz & Folk, 1984) which can influence the 13C value of

precipitated carbonate and often result in material with negative values, while travertines

forming from high temperature geothermal springs have more elevated values (Fouke, 2011;

Renaut et al., 2013; Turi, 1986). Despite the complexities, sedimentary records obtained from

such localities are sensitive to variations in water balance, input of OM and productivity. This

combined with the anoxic bottom conditions mean that sedimentary records frequently

provide excellent palaeoenvironmental archives (Schwalb et al., 1999).

Speleothems form as ground waters charged with CO2 seep into caves. Here the waters degas

CO2, forcing the precipitation of calcite. In instances where the ground waters have high Mg/Ca

ratios, aragonite can also form (Lambert & Aharon, 2011; Schwarcz, 1986). Because

speleothems have been used extensively for palaeoclimate interpretations, numerous reviews

exist covering the principals and caveats controlling the 13C and 18O values and trace element

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concentrations in speleothems (Fairchild & Baker, 2012; Hartland et al., 2012; Hendy & Wilson,

1968; Schwarcz, 1986).

The clumped isotope method has been applied to speleothems, but most studies show that

values are offset towards higher than expected temperatures (Affek et al., 2008a; Affek et al.,

2008b; Kluge & Affek, 2012). Because the formation of speleothems may be analogous to the

freshwater diagenesis of carbonates it will be important to examine well-constrained diagenetic

environments.

Hydrogen

Isotopes of H (1H and 2H) are useful in understanding carbonate diagenesis because H is

incorporated as water into fluid inclusions within carbonates and is available for comparison

with the oxygen in the same inclusion. Because H is not incorporated into the carbonate

structure to any great extent, the stable isotopes of H are not fractionated during carbonate

deposition and therefore the isotopic composition of H reflects the origin and evaporation

history of the fluid. Normally the 2H value of water, which has not been fractionated during

evaporative processes, has a defined relationship relative to the 18Ow value, namely the

meteoric water line (MWL; Craig & Gordon, 1965). Evaporated fluids deviate from the MWL

depending on the relative humidity in the ambient environment (Gonfiantini, 1986). Waters

evaporating in a humid environment fall closer to the MWL compared to those in more arid

localities. Because the magnitude of the fractionation effects is different for hydrogen and

oxygen, an evaporating solution can take different pathways depending upon the initial

chemical composition (Fig. 4).

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Boron

As a result of the long residence time of B, the oceans have a fairly homogeneous 10B and 11B

ratio, with a 11B value of ca 39.5‰ (Foster et al., 2013). This value represents a balance

between the major input of B from the continents (-8 to -15‰; Ishikawa & Nakamura, 1993)

and removal by processes which act to increase the 11B value, such as adsorption onto clays,

exchanges with the oceanic crust and incorporation into carbonates. The incorporation of B

into modern carbonates, which have 11B values of +22±3‰ (Hemming & Hanson, 1992)

represents about 20% of the loss of B from the oceans (Vengosh et al., 1991). Within natural

waters B exists complexed as either B(OH)3 or B(OH4)- depending upon the pH of the solution.

Because there is significant isotopic fractionation of 11B between these species (Zeebe, 2005b)

and because it is generally believed that only the B(OH4)- ion is incorporated into the crystal

structure of carbonates, probably substituting for the CO32- ion, the 11B values of the

carbonates become more positive as the pH increases. The promise is that the 11B value of

unaltered carbonates will provide information on both recent and ancient oceanic pH. There

has been a significant amount of work to test this hypothesis concentrating on calibrating the

11B values of the skeletal carbonate of various organisms in situations where the pH of the

environment is well-constrained (Hoenisch et al., 2008; Sanyal et al., 1996). However, it is

unclear how the 11B values of the skeletons of organisms such as corals, which have been

hypothesized to elevate the concentration of H+ at the calcification site (McConnaughey, 2003),

relates to pH variations in the external environment. There have been measurements of the

11B values of unaltered Mesozoic fossil material (Foster et al., 2010) in attempts to derive

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oceanic pH as well as in in older time periods such as the Neoproterozoic using non-skeletal

carbonates (Kasemann et al., 2010; Ohnemueller et al., 2014).

Sulphur

Sulphur has three stable isotopes, 32S, 33S and 34S, and is mainly present in the ocean as

sulphate (SO42-), a species involved in the oxidation of OM in the absence of oxygen. The

significant fractionation of 34S that occurs during this process (sulphate reduction) results in 32S

being preferentially incorporated into H2S while the residual SO42- is enriched in 34S (Chambers

& Trudinger, 1979). The fractionation accompanying this step is between 40‰ and 70‰. In

the presence of Fe, pyrite or other related iron sulphide minerals are formed. The present 34S

of the oceanic sulphate is approximately +20‰ and because SO42- is incorporated into

evaporite minerals, such as gypsum and anhydrite with only a minimal amount of fractionation

(Holser & Kaplan, 1966; Lloyd, 1968), the 34S values of evaporites can be used to monitor the

changing 34S values of the oceans through time (see later discussion). An additional proxy of S

in the oceans is recorded in the 34S values of barite (BaSO4), a mineral formed throughout the

water column by a variety of different organisms (Griffith & Paytan, 2012; Paytan et al., 1998)

and also as a diagenetic mineral formed within the interstitial pore water.

Although the S cycling is intimately linked to the carbon cycle, SO42- is also included as a trace

constituent in carbonate minerals. This S is known as carbonate associated sulphate (CAS)

(Lyons et al., 2005). Workers using this proxy have expended significant effort to make sure

that all traces of the S adhering to the exterior of the carbonate have been removed before

extracting the S and converting it into a form suitable for S-isotopic analysis (Gill et al., 2008;

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Gill et al., 2011). There is not believed to be any isotopic fractionation during this step and

therefore the34S value of this CAS is thought to reflect oceanic compositions, and

consequently has been used as an additional indicator of the 34S value of the oceans through

time (Gill et al., 2011; Paris et al., 2014). The abundance of 33S within CAS can also be measured

and its concentration has implications regarding the evolution of oxygen in the atmosphere

(see later discussion).

Magnesium

The behaviour of Mg isotopes (24Mg, 25Mg and 26Mg) during modern carbonate deposition has

been studied by Wombacher et al. (2011) who determined small, but significant, differences

between the fractionation exerted by certain aragonitic (-0.9 ± 0.2‰) and calcitic organisms

(-2.6 ± 0.3‰), with aragonite generally being less depleted than calcite. In contrast Mg-clays

are slightly enriched in 26Mg. Initial work on 26Mg reported a small temperature dependence in

both inorganic aragonite (Wang et al., 2013) and corals (Saenger et al., 2014). While limited in

extent, studies so far appear to show discrimination against the heavier isotopes during

diagenesis, so that diagenetic carbonates such as dolomites are isotopically more negative than

modern carbonates (Carder et al., 2005; Galy et al., 2002; Geske et al., 2015; Higgins & Schrag,

2010; Higgins & Schrag, 2012).The relative discrimination of Mg isotopes in carbonates and clay

minerals make them a powerful tool for studying the processes of geochemical cycling of Mg

(Higgins & Schrag, 2015).

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Calcium

Calcium has five stable isotopes (40Ca, 42Ca, 43Ca, 44Ca and 46Ca). The most common minor

isotope investigated is mass 44, reported either to mass 40 or mass 42 [44/40Ca or 44/42Ca;

mass 42 is used as a result of interferences at mass 40 from the argon used in the plasma on

the ICP–MS (inductively coupled plasma – mass spectrometer)]. The modern oceans appear to

be homogenous with respect to Ca isotopes with differences in the isotopic composition of

sources and sinks (De La Rocha & DePaolo, 2000; Zhu & Macdougall, 1998) and variations in the

rate of cycling of calcium through sedimentary rocks (Holmden, 2009; Holmden et al., 2012)

producing secular variation of about 2‰ in the 44Ca value of unaltered fossil calcareous

organisms (De La Rocha & DePaolo, 2000; Farkaš et al., 2007). The 44Ca value has been

investigated in a range of different modern calcareous organisms. Some, such as corals, show a

small amount of enrichment (0.1 to 0.6‰; Pretet et al., 2013) while others, such as planktonic

foraminifera, show more negative values (ca -1.5‰; Griffith et al., 2008). It has been suggested

that the 44Ca value of such organisms is related to the temperature of calcification (Griffith et

al., 2008; Gussone et al., 2005; Heuser et al., 2005; Immenhauser et al., 2005), with similar

relationships being observed in calcite and aragonite. However, temperature dependence does

not appear to be ubiquitous (Sime et al., 2005) and some workers have claimed that the effects

are ‘insignificant’ when they do occur (Steuber & Buhl, 2006) . Diagenetic processes have been

noted by some workers (Steuber & Buhl, 2006), the most important of which is fractionation

during precipitation of biogenic carbonates, although there is minimal fractionation of 44Ca

during dissolution and precipitation reactions (Fantle & DePaolo, 2007). This allows the 44Ca

value to be used to calculate rates of recrystallization.

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Strontium

Although strontium has three stable isotopes, 86Sr, 87Sr and 88Sr, the most common

measurement made in carbonate materials is that of 87Sr relative to 86Sr (87Sr/86Sr). Because

carbonate organisms do not fractionate the 87Sr/86Sr when they form skeletons, the 87Sr/86Sr of

the oceans can be measured through time in unaltered carbonate material (Burke et al., 1982;

DePaolo & Ingram, 1985; McArthur et al., 2001). The temporal change in the Sr-isotope ratio of

the oceans reflects the erosion of continental crust which has relatively high 87Sr/86Sr ratios

(87Sr/86Sr > 0.7140) with the 87Sr being supplemented by the decay of 87Rb and increased cycling

of ocean water through oceanic ridges which have relatively low 87Sr/86Sr values (87Sr/86Sr=

0.7040; Brass, 1976). The steady increase in 87Sr/86Sr over the past 60 Myr for example has

been linked to Himalayan uplift that has delivered an increased flux of radiogenic Sr to the

oceans (Raymo & Ruddiman, 1992). Similar arguments have been used to explain the evolution

of the 87Sr/86Sr ratio of seawater during earlier periods in Earth history (Montanez et al., 1996).

Measurements have also been made of the 88Sr/86Sr ratio in carbonates. Because this ratio is

not altered by radioactive decay, its measurement provides an additional constraint for

understanding carbonate deposition and diagenesis, as well as the Sr cycle in the oceans

(Boehm et al., 2012; Krabbenhoft et al., 2010). The 88Sr/86Sr ratio is reported relative to NIST

SRM 987 using the conventional delta notation (88Sr), with the total range of reported values

less than 0.2‰. Carbonates measured to date have a 88Sr value of ca 0.2‰ compared to

seawater at ca 0.39‰, suggesting a fractionation of about 0.2‰.

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Additional Isotopic Tracers

Additional isotopic systems which are potentially available for future research into carbonate

diagenesis include elements such as Cr (Wille et al., 2013), Fe , Cu, Zn, (Conway et al., 2013), Mo

(Wille et al., 2013) and Hg (Kritee et al., 2009; Kwon et al., 2013). The isotopic systematics of

these elements have only recently been explored and are not therefore widely applied to

carbonate systems. For example, the isotopic systematics of Mo and Cr in non-carbonate facies

have helped to define the redox state during the Archean (Crowe et al., 2013; Wille et al., 2013)

contributing towards current understanding of oxygen levels during the Precambrian.

Conceivably these indicators could also be applied to carbonate systems.

Mass Independent Isotopic Fractionation

Systems which have more than two stable isotopes can exhibit mass independent fractionation,

so that the behaviour of one of the minor isotopes is not directly predictable from the

measurement made on the second minor isotope. The classic example of mass independent

fractionation is that exhibited by the 18O, 17O and 16O system. On the Earth’s surface the

fractionation of 17O/16O can be related to that of 18O/16O through the terrestrial fractionation

line (TFL) which has a slope of 0.53 equating to the difference in the zero-point energy of the

two isotopic systems. Hence, a plot of the 17O value against the 18O value in most terrestrial

systems will produce samples which fall on a line with a slope of 0.53. Samples which are

affected by specific types of nucleosynthetic processes, such as certain meteorites, might

contain materials enriched in one of the stable isotopes and therefore fall off of the TFL.

Although early work suggested that deviations from the TFL were solely explained by

nucleosynthetic processes (Clayton et al., 1977; Clayton & Mayeda, 1977), it is now recognised

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that interactions between ultraviolet radiation and gaseous molecules such as O2 can lead to

much of the observed mass independent behaviour (Thiemens, 2006; Thiemens & Heidenreich,

1983).

The mechanisms producing different fractionation are not precisely known although there may

be several explanations, including a process known as self-shielding and/or various chemical

processes (Thiemens, 2006). During the modern day non-mass dependent fractionation is

restricted to the upper atmosphere where ultraviolet light has not been diminished by the

presence of ozone. As a result, atmospheric O2 has a small but measureable 17O anomaly.

Similar behaviour has been recognised in the sulphur isotopic system prior to the Great

Oxygenation Event (GOE), around 2.3 Ga, so the deviations from the expected behaviour of 33S

from that of 34S have been interpreted to reflect the evolution of ozone in the atmosphere

(Farquhar et al., 2000). Once sufficient oxygen was produced from the activity of

photosynthetic organisms, ozone reduced the amount of ultraviolet light reaching the Earth’s

surface, thereby inducing a mass independent isotopic effect in the sulphur system. Hence, in

the fossil record, sulphate extracted from sediments prior to the GOE frequently showed

deviations from the TFL, while, without exception, all sulphate samples analyzed in sediments

deposited after the GOE show no mass independent fractionation. How and whether mass

independent fractionation theory might be applied to diagenetic carbonate systems has not

been explored in detail. Clearly sediments formed before the GOE, might have the sulphur

isotopic anomaly removed through diagenetic processes, but the reverse is not possible. In the

case of oxygen, although the majority of meteoric waters fall on the TFL some precipitation

occurring at high latitudes and recorded in ice cores contain small anomalies (Landais et al.,

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2008; Risi et al., 2010). Anomalies can also arise as a result of mass independent fractionation

during evaporation of water (Angert et al., 2004) and such signatures might be incorporated

into carbonates. In addition, carbonates diagenetically altered prior to the GOE might have an

oxygen isotopic anomaly as precipitation during these time periods would have been strongly

influenced by ultraviolet radiation.

Minor and Trace Elements

Trace and minor elements contained within carbonate minerals can be incorporated either as a

substitute for one of the major structural groups, Ca2+ or CO32-, adsorbed onto the external

crystal surfaces, included as lattice defects and/or as contaminant mineral phases or fluid

inclusions (McIntire, 1963). In the case of Mg and Sr, only 1% of the concentration appears to

be in exchangeable sites within the carbonate (Amiel et al., 1973a) , the rest being substituted

for Ca. Although in practice all elements are incorporated into carbonates to some extent, they

can be separated into the minor elements (Sr, Mg and Na), usually present at concentrations

greater than 100 to 1000 ppm and the trace elements, usually present at concentrations of less

than 1 to 10 ppm.

In open marine environments the concentrations of trace elements in carbonates are generally

very low (<1 ppm) and it becomes difficult to separate elements which are truly substituted for

Ca in the crystal lattice from those which have absorbed onto the exterior crystal surfaces or

those which are present as contaminants. Many elements which are quite common in the

Earth’s crust, such as Fe, are only present at very low concentrations in seawater (<5 ppb) and

consequently have low concentrations in carbonates. This is because while the elements may

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be very soluble in river water, they are very redox sensitive and precipitate out very easily as

metal hydroxides in seawater. Even the soluble portion is largely believed to consist of colloids

complexed to OM (Wu & Boyle, 1998). Such oxides and colloids are deposited rapidly near river

mouths, but also travel long distances behaving very much like clay particles. Many of the

studies which report high values of elements such as Mn, Fe and others (Friedman, 1968;

Livingston, 1971) in carbonates probably reflect variable amounts of contamination. Such

contamination arises from fluvial input and/or atmospheric dust (Swart et al., 2014). The

metals are either coprecipitated with the carbonates, form coatings on the surfaces of the

carbonates or are precipitated as authigenic minerals and present within fluid inclusions. True

trace element values within the carbonate itself are probably better represented by the

concentrations of Fe, Zn and Cu (<1 ppm; St. John, 1974), and Pb and Cd (ca 10 ppb; Shen &

Boyle, 1988) in corals.

The most useful environmental information is generally obtained from elements which have

substituted for calcium. This substitution is described by a partition or distribution coefficient

(Doerner & Hoskins, 1925) which relates the ratio of the element (M) under consideration to

that of Ca in the mineral and fluid phase (Eq. :

(1)

If the value of D is one then there is no preferential accumulation of the contaminant element,

if D is greater than one the mineral preferentially incorporates the metal and if D is less than

one the metal is discriminated against. As a general rule, elements with ionic radii larger than

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Ca, have higher distribution coefficients for incorporation into aragonite and lower D values for

incorporation into LMC and HMC. During carbonate precipitation and diagenesis, smaller radii

trace elements (Fe, Mn, Zn, Cu and Cd) have distribution coefficients greater than one (Crocket

& Winchester, 1966; Pingitore, 1978) and are therefore concentrated into the diagenetic phase.

The larger radii elements (Pb, Ba and U) act more like Sr and are discriminated against

(Pingitore, 1984; Pingitore & Eastman, 1985). Although the range of concentrations and

distribution coefficients for elements incorporated into the three common forms of calcium

carbonate (aragonite, LMC and HMC) have been tabulated by Veizer (1983), there are wide

ranges in the estimates of the distribution coefficient for most trace and minor elements,

making their practical application difficult.

Such uncertainty arises from the differences in the analytical procedures used in determining

the D values, as well as an absence of research in this area. For example, although Veizer

(1983) reported a range of DMn values of between 5 and 30, it is clear that this range is

controlled by the kinetics of the precipitation with the low values being manifested at very

rapid rates and the high values at very slow rates (Lorens, 1981). Although similar inverse

correlations between precipitation rate and D values were reported for Co and Cd, it is likely

that other small radii elements have similar relationships.

In contrast to the smaller radii elements, the DSr for calcite showed a positive correlation with

precipitation rate (Lorens, 1981). Strontium has a distribution coefficient into aragonite of

approximately one yielding concentrations of ca 7000 ppm in the mineral precipitated from

seawater (Banner, 1995; Kinsman, 1969). However, the distribution coefficient is temperature

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dependent, decreasing with increasing temperature in both inorganic and organic precipitates

(Beck et al., 1992; Kinsman & Holland, 1969; Smith et al., 1979). It has also been suggested that

the D value is significantly higher in inorganically precipitated aragonites suggesting that the Sr

concentration could be used to distinguish between organically and inorganically produced

aragonite (Milliman et al., 1993). Some biogenic aragonites (i.e. Pteropods) have low Sr

concentrations (ca 1000 ppm) suggesting that organisms can exercise control on the mineralogy

and chemistry of their skeletons. In particular, organisms with evolutionary advanced

calcification mechanisms can exclude certain ions. Biogenic LMC organisms typically have Sr

concentrations between 1000 and 1200 ppm (corresponding to a DSr of ca 0.1), a value close to

that measured in inorganically precipitated LMC (Holland et al., 1964). Values for the

alteration of aragonite to LMC or biogenic LMC to inorganic LMC have been measured to be ca

0.05 (Baker et al., 1982; Katz et al., 1972a). Such differences are related to the rate of

dissolution of aragonite and consequent precipitation of calcite compared to the direct

precipitation of calcite. The values correspond extremely closely to the data from Lorens

(1981) who found the DSr for LMC ranged from values close to 0.1 [similar to the values

determined by Kinsman (1969)] at high rates to very low values at low precipitation rates

[<0.05; similar to but even less than the values reported by Baker et al. (1982) and Katz et al.

(1972b)]. High-Mg calcites have higher Sr concentrations than LMC, perhaps as a result of

lattice distortion (3000 to 4000 ppm; Milliman, 1974).

Barium

Barium in carbonates is believed to be related to the Ba/Ca ratio in seawater (Lea et al., 1989)

and, because of its nutrient type behaviour (Lea & Boyle, 1991), has been used as a proxy

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indicator of upwelling in foraminifera (ca 0.3 ppm) and corals. Elevated Ba concentrations in

coral skeletons, which normally have concentrations of ca 2 to 3 ppm Ba, have also been used

as a tracer of riverine runoff (Alibert et al., 2003; McCulloch et al., 2003) although non-fluvial

signals have been observed in some corals (Sinclair, 2005). Similar riverine influences on Ba

concentration have been observed in foraminifera (Hoenisch et al., 2011).

Lead

Lead is present in low concentrations in all modern marine carbonates. Lead levels are

sensitive to variations in the external environment, and changes as a result of the addition of Pb

to petrol can be seen clearly in many carbonate records such as corals and sclerosponges

(Rosenheim et al., 2006; Shen & Boyle, 1988).

Uranium

Uranium shows relatively high concentrations (2 to 8 ppm) in aragonite, but low values in LMC

(0.02 ppm) and HMC (Amiel et al., 1973b; Chung & Swart, 1990; Gvirtzman et al., 1973; Schoepf

et al., 2014; Schroeder et al., 1970). Because U is mainly complexed with carbonate ions in

seawater, it has been suggested as a potential proxy of the pCO2 in the oceans (Swart &

Hubbard, 1982).

Rare Earth Elements

The rare earth elements (REE) refer to a series of 15 elements of increasing atomic weight and

decreasing atomic radii (1.14 Å for La to 0.84 Å for Lu) which generally occur in the same

mineral deposits and have similar geochemical properties. The REE distribution or pattern

usually refers to the REE concentration normalized to a standard reference material such as

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chondritic meteorites or certain shales [for example North Atlantic Shale Composition (NASC)].

Modern carbonates incorporate REEs in ratios similar to those in seawater (Sholkovitz & Shen,

1995) and the REE signal or REY (REE with yttrium) has been used as a proxy for runoff

(Wyndham et al., 2004), pollution (Fallon et al., 2002) and bioproductivity (Wyndham et al.,

2004). Certain REEs such as Ce and Eu exist in multiple valence states. Cerium which has both

Ce3+ and Ce4+ states is mainly present in seawater as the Ce4+ form and is relatively insoluble.

Therefore carbonates formed from seawater tend to have a reduced Ce concentration or a Ce

anomaly in the normalized REE pattern. Europium also has multiple valence states and can be

present as either Eu2+ or Eu3+. Because Eu has an ionic radius similar to Ca2+, higher

concentrations of Eu tend to be present in carbonates formed in reducing environments.

Non-conventional Minor and Trace Elements

Another category of trace elements, including univalent elements such as Na+, K+ and Cl-, as

well as anionic complexes such as SO42-, PO43-, NO3- and B(OH)-, are also incorporated in

carbonates (Ishikawa & Ichikuni, 1984; Land & Hoops, 1973; Lyons et al., 2005; Prokopenko et

al., 2013; Staudt et al., 1993; Staudt & Schoonen, 1995; Veizer et al., 1977; White, 1977). These

are termed non-conventional indicators (NCIs) because it is not known precisely where or how

they become incorporated in the mineral structure. Univalent cations might ‘fit’ into the same

locations in the crystal lattice as Ca2+, but there is a problem of charge balance. Where anions

might end up is a matter of speculation. For example, complexes such as SO42- or PO43 might

substitute for a CO32- group in the carbonate structure. Undoubtedly, a large proportion of the

total concentration of these species is either adsorbed onto the exterior of the crystal surface,

is present as contamination or is contained within fluid inclusions. The main use of variations in

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the concentration of Na, K and Cl has been as an indicator of salinity, basically the higher the

concentration of the element, the higher the salinity value (Ishikawa & Ichikuni, 1984; Land &

Hoops, 1973; Veizer et al., 1977; White, 1977). However, as a result of contamination issues,

the equilibrium concentrations of these elements have not yet been defined. Variations in the

concentration and stable isotopes (N and S) of P, N and S are suggested to be related directly to

the concentration of these elements in the environment (Lyons et al., 2005; Montagna et al.,

2006). However, the concentrations of all of these constituents are easily influenced by

contamination and prior to analysis samples must be rigorously cleaned. Even when such

cleaning takes place it is uncertain whether all contamination, or in some instances too much,

has been removed. In addition, microbial processes taking place within the carbonate have

been shown to drastically alter the concentration and isotopic composition of some of these

proxies. For example NO3- is present in the open oceans at very low concentrations (<0.1 M),

yet within skeletons organic material can degrade to NH4+ and then NO3- substantially

increasing its concentration within the interstices. The interior of carbonate particle may

become anoxic leading to denitrification and thereby causing the 15N value to become

elevated. Because during such processes carbonate dissolution and precipitation also takes

place, the NO3- produced and incorporated into the carbonate will reflect diagenetic rather

than original conditions (Diaz et al., 2013) .

Effect of Temperature

Temperature frequently exerts a control on the minor element composition of carbonate

minerals. In the case of DSr in aragonite there is a strong inverse correlation between

temperature and Sr content in both inorganic (Kinsman & Holland, 1969) and biogenic systems

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(Weber, 1973). The slopes of the temperature dependence are different for biogenic and non-

biogenic systems with the slope for biogenic aragonite, for example corals, being about twice as

steep as that in inorganic aragonite. Although less information exists for LMC, the experiments

of Baker et al. (1982) and Katz et al. (1972a) showed that the DSr was not influenced by

temperature in inorganically formed LMC. Perhaps a failure to observe an expected decrease

with increasing temperature can be explained by a kinetic control on the DSr for calcite, with the

DSr increasing at higher temperatures (Lorens, 1981). Higher rates of precipitation cause higher

DSr values thus offsetting the thermodynamic tendency for DSr to decrease at higher

temperatures. The distribution of other smaller radii elements (Mn, Co and Cd) tends to

decrease with increasing rates of precipitation (Lorens, 1981).

In contrast to Sr, the distribution coefficient of elements such as Mg in biogenic calcite (Chave,

1954) increases with temperature. This property has led to widespread use of the Mg/Ca ratio

as an indicator of temperature in organisms such as foraminifera (Lea, 2002; Lea et al., 1999;

Lear et al., 2000). Once the temperature has been obtained using the Mg/Ca ratio, it can be

combined with the 18O value of the same organism to determine the 18Ow. Salinity can then

be estimated using a published relationship between the 18Ow value and salinity (Flower et

al., 2004; Schmidt et al., 2004). However, while the Mg/Ca ratio is related to temperature,

properties such as salinity, carbonate ion content and pH also appear to influence the ratio

(Arbuszewski et al., 2010), thereby confusing the relationship at least in the tests of

foraminifera. Calculation of salinity using this method is also hampered by large errors

associated with the equation relating the Mg/Ca ratio and 18O value to temperature as well as

uncertainties regarding the relationship between salinity and the 18Ow value.

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A number of other elements also show temperature dependence in biogenic aragonite

including U (Min et al., 1995; Shen & Dunbar, 1995), Mg (Mitsuguchi et al., 1996) and Ba

(Gaetani & Cohen, 2006). However, because there are frequently a number of other controls,

the relationships with temperature have proven to be variable. For example, in the study of

Fallon et al. (2003) a large amount of variability was determined in the relationships between

temperature and Sr, Mg, U and B in a number of corals from a transect stretching from the

inshore to the outer barrier. It has also been proposed that a combination of Sr, Mg and Ba

could be used to arrive at a more reliable coral thermometer (Gaetani et al., 2011). Because

coral shows different amounts of discrimination against these elements, it has been argued that

a Rayleigh based approach can be used to make accurate and precise determination of sea-

surface temperature (SST) without the need for species specific calibrations.

Secondary Minerals within Skeletons

In some instances the precipitation of secondary minerals within the interstices of ‘live’ skeletal

materials can contribute significantly to the trace element concentration of carbonates. For

example the mineral brucite (Mg(OH2) has been recorded within coral skeletons (Buster &

Holmes, 2006; Schmalz, 1965) and some calcareous algae (Weber & Kaufman, 1965) thereby

increasing the Mg content. Other secondary minerals within skeletons, such as LMC and

aragonite, have also been widely documented and are known to alter primary elemental and

stable isotopic compositions affecting palaeoenvironmental interpretations (Hendy et al., 2007;

McGregor & Gagan, 2003). Recently, the presence of a carbonate phase containing

concentrations of Mg approaching that found in dolomite has been noted in the skeletons of

some calcareous red algae with skeletons composed of HMC (Nash et al., 2013). Finally,

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skeletons such as corals and molluscs are frequently bored by endolithic organisms including

algae, sponges and fungi (Golubic et al., 1975; Perkins, 1977; Perkins & Sentas, 1976). These

remove primary skeletal material and the resultant boring frequently becomes filled with fine

mud and/or secondary cements (Bathurst, 1966). Such borings, typically most noticeable

around the exterior of the carbonate grain or skeleton as micritic rinds (Bathurst, 1966), add to

the already complicated geochemistry of the original components.

Scales of Geochemical Variation

Diagenesis affects the individual components of the rock differentially and therefore different

geochemical results can be obtained by analyzing rocks at different scales. For example,

cements often form between grains without affecting the chemical composition of the grains

themselves, or different components may dissolve and reprecipitate preferentially. In addition

crystals forming from the same solution may exhibit different trace element and isotopic

compositions in different growth sectors (Dickson, 1991; Reeder & Grams, 1987; Reeder &

Paquette, 1989; TenHave & Heijnen, 1985) (compositional sector zoning). Bulk stable isotopic

measurements, which are common in many studies, may not capture the chemical signatures

of diagenetic processes unless the entire rock has been altered. The microsampling of

carbonates has progressed from the initial attempts of Dickson & Coleman (1980) who used a

scalpel to excise sufficient material from thin sections, to more modern use of computer-

controlled microdrilling methods. However, these methods are still rather crude and do not

offer the precision necessary to sample material at spatial resolutions of less than ca 100 to 200

m. In addition, there have been attempts to use laser ablation coupled to mass spectrometers

(Dickson, 1991; Dickson et al., 1991; Larson & Longstaffe, 2007; Sharp & Cerling, 1996; Smalley

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et al., 1992) and secondary ion mass spectrometry (SIMS; Kozdon et al., 2009; Rollion-Bard et

al., 2003; Treble et al., 2005; Treble et al., 2007) to measure stable isotopes. These latter two

methods offer promise, but are still not used widely being limited by the rarity of the

equipment or poor precision of the results. More widely available methods for microsampling

minor and trace elements in carbonates include the electron microprobe, which typically can

analyze materials at concentrations down to ca 500 to 1000 ppm (Moberly, 1968), secondary

ion mass spectrometry, which can analyze at the ppm level (Swart, 1990; Veizer et al., 1987)

and measure stable C and O isotopes (Vetter et al., 2013; Williford et al., 2013), or nuclear track

methods which can analyze spatial distribution of elements such as U using fission tracks

(Chung & Swart, 1990; Swart, 1988; Swart & Hubbard, 1982) or B and Li using alpha tracks. As

stated by Hudson (1977): “…we must take our limestone to pieces” if we want to adequately

understand the diagenesis of the rock.

Fluid Inclusions

Fluid inclusions generally represent water and gas trapped during the formation of crystals

which have remained fluid during cooling to normal temperatures (Roedder, 1984). Inclusions

range in size from <2 m to >1 mm and may include water with associated salts as well as

vapour, hydrocarbons, precipitated phases and organic matter. Inclusions can be classified as

being one-phase all-liquid, one-phase all-gas, two-phase or even three-phase. Inclusions can be

primary (i.e. forming at the same time the mineral was precipitated) or secondary, associated

with alteration after the crystal growth. Fluids entrapped at low temperature should remain a

single-phase inclusion while those formed at temperatures higher than 50oC are two phase at

surface temperatures and pressures (Roedder, 1984). In addition, primary inclusions can be

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modified during secondary processes. Criteria for describing primary, secondary and

diagenetically modified inclusions are well-established (Goldstein, 2003) although

interpretations can be controversial. Assuming that primary fluid inclusions represent samples

of the formation fluid and that the inclusions have remained unaltered, information can be

obtained regarding the temperature and chemical composition of the mineral using a variety of

different analytical techniques.

The simplest technique is to examine a doubly polished wafer of rock with fluid inclusions in a

heating and cooling stage attached to a petrographic microscope. As a normal two phase fluid

inclusion (liquid and vapour) is heated, the temperature at which the two phases become

homogenized (Th) is indicative of the minimum temperature of formation (uncorrected for

pressure). Conversely when a sample is frozen and then warmed, the temperature at which the

solid phase (‘ice’) melts is related to the density of the fluids in the inclusion (Potter et al.,

1978). Although the salinity of the fluid can be approximated using a variety of assumptions, it

is important to note that the ions present are not likely to have behaved conservatively and

therefore the estimated salinity may be incorrect. Such temperature and salinity data can be

combined with the 18O value of the carbonate to calculate the 18Ow (assuming a defined

relationship between salinity and 18Ow). Ideally such information should conform to the

homogenization temperature. Differences between the temperatures obtained from fluid

inclusions and the temperature derived from the 18O value (using the salinity estimate

obtained from freezing) are usually ascribed to problems associated with the origin of the fluid

inclusions themselves, unknown pressure corrections, or problems relating salinity to the 18Ow

value (Moore, 1989; Moore & Druckman, 1981; Prezbindowski, 1987).

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The chemical composition of the fluids can be analyzed using a variety of methods [see Roedder

(1990)]. These methods have been utilized both to determine the fluids involved in diagenetic

processes and also within evaporite minerals to determine ratios of ions in evaporative fluids.

Based on these ratios, inferences have been made regarding the chemical composition of

seawater through time (Demicco et al., 2005; Hardie, 2003; Horita et al., 2002). Assuming that

the 18O and 2H values of the fluid inclusion represent the formation fluid, then the 18O value

of the carbonate can be used to estimate the temperature of formation or the evolution of the

18Ow value and the chemistry of seawater through time (Knauth & Beeunas, 1986).

Calcite and Aragonite Seas

It has been suggested that the predominant form of inorganic marine cementation has changed

from its current favoured form, aragonite, to LMC and then back again to aragonite several

times over the past 600 Myr. This ‘aragonite-calcite’ seas phenomenon, as it is known, was

documented first in the inferred mineralogy of ooids (Sandberg, 1983) and subsequently

recognised in the form of mineralogy precipitated by a number of different organisms

considered to have weak control on their calcification processes (Stanley, 2006). It has also

been suggested that some organisms are able to change the mineralogy of their skeletons (Ries

et al., 2006) depending upon the chemistry of the oceans, specifically the ratio of Mg/Ca in the

precipitating fluid. During intervals when the Mg/Ca ratio of seawater was similar to the

modern day, precipitation of aragonite was favoured while during periods of low Mg/Ca ratios

calcite was the preferred form. This notion is supported by inhibited precipitation of LMC in

solutions with high Mg/Ca ratios, the preservation of secular variations in the Mg/Ca ratio of

fluid inclusions from primary halite (Lowenstein et al., 2003), variation in the mineralogy of late

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stage evaporites supporting a change in the original Mg/Ca ratio of seawater (Hardie, 1996),

and secular variations in the Mg/Ca ratio of molluscs (Steuber & Rauch, 2005) and echinoderms

(Dickson, 2004). The Mg/Ca ratio in the oceans responds to a combination of changes in the

concentration of both elements, which in turn are thought to be related to changes in sea-floor

spreading and the circulation of seawater from mid-oceanic ridges (Demicco et al., 2005;

Hardie, 2003). In addition the change in Mg/Ca ratio is believed to be accompanied by small

changes in the 26Mg (Higgins & Schrag, 2015) and 44Ca values of seawater (Gothmann et al.,

In Review). As a consequence of the change in the Mg/Ca ratio of seawater, the nature of early

marine cement should change through time as should the 18O, 13C values and minor and trace

element composition. In addition, modification of the Mg/Ca ratio within marine pore fluids,

even close to the sea floor, can result in precipitation of marine cements which are LMC or HMC

rather than aragonite (Aissaoui, 1988; Schroeder & Purser, 1986).

Influence of Mg/Ca on Cement Type

For many years it was believed that crystal mineralogy and habit indicated whether cements

formed either in marine or meteoric environments (Folk, 1974; James & Choquette, 1983;

James & Choquette, 1984; Longman, 1980). Acicular aragonite crystals were considered

characteristic of marine dominated fluids with high Mg/Ca ratios, while LMC equant spar and

overgrowths were associated with freshwater. However, several studies have documented the

occurrence of supposed freshwater cements and fabrics in environments which could only be

influenced by marine fluids (Melim et al., 1995; Schlager & James, 1978). Because the

formation of LMC is determined by the Mg/Ca ratio, it is probable that the LMC cements

observed by previous workers in environments obviously dominated by saline waters were

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influenced by processes acting to reduce the Mg content of the pore fluids (Hendry et al.,

1996). An additional control on mineralogy can be exerted by changes in the saturation state of

the oceans, related to circulation and depth. For example, the formation of dolomite occurs in

many marine settings and reduces the Mg2+ content of the pore fluids thereby allowing LMC to

be precipitated, while in deeper waters which are undersaturated with respect to aragonite

LMC would be formed regardless of the Mg/Ca ratio of seawater.

Sea Level

Earth history is marked by periods when there were only minor changes in sea level (10 to 20

m) while during others changes of 100 m or more occurred over 20 to 100 kyr time scales.

Because the principal cause of large sea-level change is glaciation, warmer time periods in

which the amount of continental ice was reduced, coincided with more stable sea level and

temperatures which were generally warmer than at present. These greenhouse and icehouse

periods tend to coincide with calcite and aragonite seas, respectively. Hence, Cretaceous

carbonates were mainly formed of LMC without the large glacial-eustatic sea-level changes

which characterized the Pleistocene. The absence of such sea-level changes would have limited

the exposure of carbonates to fresh-water and associated diagenetic alteration. Furthermore

the presence of a primary mineralogy of calcite, rather than aragonite, means that the

geochemical signatures of carbonates formed during calcite seas are more likely to be

preserved than those formed during other times.

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Dolomitisation

The use of geochemical tracers to study dolomitisation has been reviewed by various workers

(Budd, 1997; Machel & Mountjoy, 1986; Warren, 2000) since the seminal paper of Land (1980).

Although more than 30 years have passed since the Land paper, and many important

geochemical studies have been published, knowledge of the factors controlling geochemistry of

dolomite has not changed significantly.

Oxygen Isotopes in Dolomites

Dolomites have equilibrium 18O values which are believed to be about 3 to 6‰ heavier than

LMC formed under the same conditions at 25oC. The 3‰ uncertainty arises because there are

at present nine different equations linking the fractionation between the 18Ow and the 18O

value of the dolomite () to temperature (Clayton et al., 1968; Fritz & Smith, 1970; Horita,

2014; Katz & Matthews, 1977; Land, 1983; Northrop & Clayton, 1966; O'Neil & Epstein, 1966 ;

Sheppard & Schwarcz, 1970; Vasconcelos et al., 2005; Zheng, 1999). Some of these equations

were derived experimentally at high temperatures (>200oC) and then extrapolated to lower

temperatures; others are based on the observations of naturally occurring high temperature

dolomites and extrapolating these data to lower temperatures (Sheppard & Schwarcz, 1970). In

contrast, the work of Vasconcelos et al. (2005) determined 18O values in microbial induced

dolomite formed at temperatures between 25oC and 50oC, while the experiments of Fritz &

Smith (1970) generated dolomites between 50oC and 70oC. The study of Zheng (1999) was

based on theoretical considerations. Although the Vasconcelos et al. (2005) equation should

seemingly be the most appropriate at low temperatures, the dolomites produced in those

experiments possessed very poor ordering peaks (or none at some of the temperatures), were

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calcian and might be better described as protodolomite (Graf & Goldsmith, 1956) or very high

Mg calcite (VHMC; Sibley et al., 1994). Because variations in stoichiometry and ordering can

produce artifacts and lead to variability in 18O values unrelated to environmental conditions

(Land, 1980; Vahrenkamp & Swart, 1994), these low-temperature equations might not be

appropriate for the interpretation of temperature and fluid composition in more mature types

of dolomite.

Carbon Isotopes in Dolomites

The 13C value of dolomite is thought to be slightly more positive (ca 1‰) than co-occurring

LMC (Sheppard & Schwarcz, 1970), but is in most instances similar to the sediment from which

it formed. The exception to this occurs where dolomitisation takes place in carbonate poor

sediments and/or the system is heavily influenced by the oxidation of OM (Baker & Burns,

1985; Compton & Siever, 1986; Irwin et al., 1977; Kelts & McKenzie, 1982). In spite of the slight

difference in the 13C value between LMC and dolomite, the processes affecting the 13C values

of dolomites are similar to those affecting other diagenetic carbonates. However, since

dolomite is believed to be preferentially formed in environments where degradation of OM is

abundant, dolomites from these environments may frequently have negative 13C values (Kelts

& McKenzie, 1982) if there are relatively low concentrations of carbonate materials in the

system. For example, dolomitisation is seen in areas where abundant upwelling promotes high

productivity generating siliceous sediments rich in OM (Compton & Siever, 1986; Malone et al.,

1994). Degradation of OM produces dissolved inorganic carbon in pore water with extremely

negative 13C, high alkalinity and low concentrations of sulphate. All of these factors have been

suggested to promote the formation of dolomite (Baker & Kastner, 1981).

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Minor and Trace elements in Dolomites

Dolomite exhibits a broad range in crystal ordering and composition ranging from an excess of

10% Ca relative to Mg, to a composition in which there are equal molar amounts. Similar to

LMC, the principal trace elements of interest in dolomites are Sr, Mn and Fe, and only a few

studies have critically examined other trace elements and REEs in dolomites (Shaw &

Wasserburg, 1985; Swart, 1988). The concentration of Sr in dolomite is generally lower than in

calcite formed under the same conditions as Sr is preferentially substituted into the Ca sites in

the crystal lattice (Kretz, 1982; McIntire, 1963); this predicts that Ca-rich dolomites will have

higher Sr concentrations than stoichiometric dolomites and was shown to be applicable to

dolomites from a wide range of ages (Vahrenkamp & Swart, 1990). Other elements with larger

radii (Ba and Pb) should show similar patterns to Sr, while elements with smaller radii (Ni, Cu,

Co, Zn, Mn and Fe) should be elevated in stoichiometric dolomites. While this may be the case,

natural variability and input of Mn and Fe from sources unrelated to the control of the Mg/Ca

ratio in dolomites probably overwhelms the stoichiometric control (Vahrenkamp & Swart,

1994). In the case of Mn, whether it is substituted within the Mg or Ca sites, can be measured

using electron spin resonance (ESR; Wildeman, 1970). This and later studies (Lumsden & Lloyd,

1984) indicate that there is more Mn in the Mg sites (compared to Ca sites) by factors of

between ca 1 to 70, a considerably greater range than the value of 1.5 presented in the model

of Kretz (1982). Lumsden & Lloyd (1988) classified the ESR spectra obtained from dolomites

into three categories depending upon whether the position of Mn could be resolved or not.

Broadly speaking these authors determined that stoichiometric dolomites tended to have the

best definable peaks and the highest proportion of Mn in the Mg sites.

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Although not well-documented in the literature, it is known from reservoir studies using

spectral gamma ray methods, that dolomitised reservoirs are frequently elevated in uranium

relative to co-occurring limestone (Swart, 1988). A possible explanation for this is that modern

aragonite dominated carbonate sediments are elevated in U (Chung & Swart, 1990), and if

these sediments are dolomitised early then the U concentration, which is principally complexed

with the carbonate, is preserved. In contrast, if the sediments are altered to LMC first in an

open system, then U is lost and any subsequent dolomitisation is likely to produce dolomites

low in U.

As stoichiometric dolomite incorporates equal amounts of Ca and Mg, attention has been paid

to the fractionation of Mg isotopes during dolomitisation (Higgins & Schrag, 2010; Higgins &

Schrag, 2012). The principal effect appears to be the formation of dolomite in which the lighter

more abundant isotope is preferentially incorporated, leading to more negative 26Mg values in

the dolomites and positive values in the pore water.

Strontium Isotopes

The 87Sr/86Sr ratio of dolomites has been used to help constrain the timing of dolomitisation

provided the depositional age is known and the sources of Sr constrained (Saller, 1984). In

situations where the only possible sources of Sr are provided by the original sediments and the

dolomitising fluids are seawater of a later age, the oldest possible age of dolomitisation can be

determined (Swart et al., 1987b; Vahrenkamp et al., 1991).

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Sulphur Isotopes

The 34S of CAS in dolomites can be used in conjunction with the concentration of NCIs to

elucidate the environment of dolomitisation. For example, because it is well known that many

dolomites are formed within the sulphate reduction zone, these should have lower

concentrations of S, normal concentrations of Na, K and Cl, but slightly elevated 34S values.

Dolomites associated with evaporite minerals might have low concentrations of SO42- (as SO42- is

removed during the formation of evaporite minerals), normal 34S values, but elevated values

of Na+, K+ and Cl-. Dolomites formed from brines which have not attained saturation with

respect to gypsum or anhydrite might be expected to show elevated concentrations of all NCIs,

including sulphate, and normal 34S values.

Clumped Isotopes and Dolomite

The clumped isotope technique should be able to ascertain which of the various equations

which link the 18O value of dolomite to temperature and the 18Ow value is correct. While a

number of studies have measured the 47 in dolomites (Ferry et al., 2011; Loyd et al., 2012),

reporting both temperatures and 18Ow values, most of these have used a theoretical equation

(Guo et al., 2009) to calculate the temperature from 47 and the 18O–18Ow to temperature

equation of Vasconcelos et al. (2005) to calculate the 18Ow value. Other studies have used

alternative 47 temperature equations, typically the original Ghosh et al. (2006) equation

adjusted for the absolute reference framework (Dennis et al., 2011) or the Passey & Henkes

(2012) equation (Geske et al., 2015; Murray et al., 2014; Sena et al., 2014). These equations are

combined with a range of different 18O–18Ow to temperature equations. A recent study also

suggests that the acid fractionation of 47 at 90oC for dolomite may be greater than the value of

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0.09‰ reported for calcite by Henkes et al. (2013) (Murray et al., 2013; Murray et al., 2014).

The use of this higher fractionation factor combined with different 47- temperature and 18O–

18Ow temperature equations produces a range of temperatures and 18Ow values which makes

precise application of the results for the moment problematic.

Models for Early Dolomite Formation

The formation of dolomite requires a source of Mg2+ (usually seawater), a mechanism of Mg2+

supply and a sediment body with sufficient permeability. Although seawater is approximately

1000 times oversaturated with respect to dolomite, any process which raises the saturation

state of dolomite has the potential to increase the efficiency and rate of dolomitisation. Such

processes would include elevation of the salinity of the fluid and/or an increase in alkalinity

caused by oxidation of OM. Clues regarding the interpretation of the geochemical signals left

behind during dolomitisation can be found in recently formed material from areas such as Abu

Dhabi and the Bahamas where dolomites have formed under reasonably well-constrained

conditions. In Abu Dhabi, Holocene dolomites are unambiguously associated with the

evaporation of seawater and the formation of evaporites (Patterson & Kinsman, 1982), and in

some places the formation of microbialites (Bontognali et al., 2012). Both the waters and

dolomites associated with the evaporites have elevated 18O values (18Ow = +2 to +6‰ and

18O =+2 to +3‰; McKenzie, 1981; McKenzie et al., 1980). However, curiously the dolomites

from the modern evaporites in Abu Dhabi are not that much more elevated in 18O than the

Plio-Pleistocene Bahamian dolomites (+3 to +4‰), where there is no evidence of evaporite

deposition (Dawans & Swart, 1988; Fig. 8) and which are generally considered to have formed

from normal seawater with a 18O value of about +1‰. It is possible that the absence of higher

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18O values in the Abu Dhabi dolomites reflects the fact that dolomites formed at low

temperatures are initially very poorly ordered and are Ca-rich. Such dolomites might initially

have more negative 18O values than well-ordered and stoichiometric dolomites eventually

recrystallizing to become more ordered and stoichiometric (Vahrenkamp et al., 1991). The

similarity of the 18O value of the dolomites from the Bahamas, which formed from normal

seawater, and those from Abu Dhabi, which are hypersaline, emphasizes the uncertainty in the

use of 18O values to ascertain the nature of the environment of dolomitisation. Older

dolomites and those formed at greater depths tend to have more negative 18O values

reflecting formation or recrystallisation under higher temperatures. One possible model for

dolomite formation maybe that the mineral starts as disordered dolomite or VHMC and

gradually recrystallises to more ordered and stoichiometric forms. In this way the Mg

necessary for dolomite formation is provided early, but the final mature mineral is formed

much later. Along this pathway geochemical patterns, such as stable isotopes and trace

elements, change leading to a chemical composition unrelated to the original environment.

Some indication of this can be observed in studies which show an increase in ordering with

depth (Gregg et al., 1992) although more convincing evidence of the transformation of early

protodolomite to mature dolomites is needed (Gregg et al., 2015).

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Behaviour of Stable Isotopes and Trace Elements during Diagenesis

Surface Related Processes

Freshwater

When sea-level is lowered, shallow water carbonates are exposed to the influence of meteoric

fluids and a series of recognizable diagenetic zones are formed (Fig. 9). Such zones include

those associated with sub-aerial exposure surfaces, vadose and freshwater diagenesis, the

mixing-zone and the marine phreatic zone. As sea-level oscillates, the partially-lithified

sediments and rocks are sequentially affected by each of these diagenetic environments,

continually overprinting primary carbonates and already diagenetically altered rocks.

Therefore, the changes observed within a carbonate sequence exposed to freshwater

conditions are usually a composite of a large number of superimposed physical and chemical

diagenetic changes. However, it is likely that the first diagenetic episode will be the most

influential and that subsequent episodes will not affect material already stabilized, but rather

affect only those remaining carbonates not yet influenced by the initial diagenetic event.

Mineralogy

Rainwater contains only small quantities of dissolved salts, derived from aerosols and

atmospheric CO2. The high CO2 and low concentration of Ca2+ causes this water to be corrosive

to all carbonate minerals. Once in the vadose zone, the rainwater (now groundwater) acquires

additional CO2 from the decay and respiration of local OM and initially dissolves the local

calcium carbonate until saturation is attained with respect to the ambient carbonate

mineralogy. In geologically young terrains, the carbonate sediments are mainly composed of

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aragonite and HMC. Because these minerals are less stable than LMC, precipitation of LMC can

occur before dissolution of aragonite and HMC is complete (Budd, 1988). Precipitation of LMC

causes further undersaturation with respect to aragonite and HMC and the system could

theoretically dissolve the metastable minerals and precipitate the more stable ones until all the

aragonite and HMC are consumed. Further additions of freshwater to the system will then

result in dissolution of LMC, at least in the upper portion of the aquifer.

Stable Carbon and Oxygen Isotopes

Although carbonate cementation in limestones was believed by many to result primarily from

exposure to freshwater (Shinn, 1969; Shinn, 2013), Clayton & Degens (1959) suggested that

13C values could help to distinguish between freshwater and marine limestones. The first

studies to utilize 13C and 18O values within diagenetic carbonates were published by Gross

(1964) using samples from the island of Bermuda. These studies showed that the 18O values of

secondary carbonates were controlled primarily by the 18O value of rainfall on the island, while

the 13C value was governed by CO2 derived from the soil zone. Gross (1964) concluded that:

“The 18O/16O and 13C/12C ratios of the limestone appear to be useful in determining their

diagenetic history where the geologic relationships are relatively simple and when the original

isotopic composition of the carbonates and water is known.”. The recognition that diagenetic

carbonates acquire the 18O signature of their recrystallizing fluids is the basis for numerous

other interpretations of the behaviour of 18O values in diagenetic carbonates. However, since

the 18O value of the carbonate also depends on the temperature of alteration there has always

been ambiguity in the interpretation of carbonate 18O values unless an independent

temperature estimate, such as that obtained from fluid inclusions or clumped isotopes, has

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been available. An example of this dilemma is shown in Fig. 3 which shows lines representing

the relationship between temperature and the 18O values of calcite or dolomite (see below for

discussion of dolomite) for fluids with varying 18O values. Plotting a data set on such a

diagram necessitates knowledge of the 18O value of the solid as well as either the original

temperatures of deposition, or the 18Ow. Ignoring for the moment the problem of

distinguishing variations in the 18O value of the carbonate caused by temperature as opposed

to the 18Ow, most of the work dealing with early diagenesis has interpreted the changes in the

18O value of the carbonates as originating from the fluids. This is because most early

diagenetic changes are considered to take place at temperatures between 20oC and 30oC which

would introduce an uncertainty of only a few per mille in the 18O value. As a consequence,

variations in the 18Ow values are considered to dominate the early diagenesis of carbonates.

Trace and Minor Elements

Changes also occur in the composition of trace and minor elements during diagenesis (Budd &

Land, 1990; Land & Epstein, 1970; Saller & Moore, 1991). Such changes usually involve the

dissolution of metastable carbonates (aragonite and HMC) releasing trace elements into the

pore water, followed by precipitation of more stable phases such as LMC and dolomite.

Variations in the concentrations of Mn and Fe in the pore fluids or varying rates of precipitation

can lead to the formation of cements with luminescent characteristics (Frank et al., 1996;

Meyers, 1974), characteristics which have been used to establish paragenetic sequences and

identify freshwater diagenesis. It is believed that high concentrations of Mn (> ca 50 ppm) lead

to luminescence, but this can be altered in the presence of Fe which tends to quench the

phenomenon. However, some workers have induced changes in luminescence without altering

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the chemical composition of the fluid (Have & Heijnen, 1985), while others have implicated

other activator elements (Machel, 1987a) including REEs.

Rare Earth Elements

The diagenetic behaviour of REEs has been investigated in a few studies (Johannesson, 2012;

Scherer & Seitz, 1980; Shaw & Wasserburg, 1985; Webb et al., 2012; Webb et al., 2009). The

most recent study shows that meteorically altered carbonates appear to have similar values

when compared to the original sedimentary components (Webb et al., 2009) although a Ce

anomaly was noted together with a slight depletion in the light REEs. However, ubiquitous

contamination from oxide phases and diagenetic incorporation from pore waters confound the

use of REEs (Della Porta et al., 2015; Haley et al., 2004; Reynard et al., 1999). Laser ablation

methods, coupled to ICP-MS, eliminates the ubiquitous contamination seen in bulk samples and

allows the REE signatures of individual cements to be targeted. Hence characteristic primary

signatures can be confirmed in cements believed to be of marine origin (Della Porta et al.,

2015).

Uranium

Uranium generally follows the behaviour expected of a large radii element and freshwater

cements generally have much lower concentrations compared to the original biogenic

components. However, it has been noted that U concentrations with freshwater phreatic

cements are significantly higher than in vadose cements, a phenomenon which was ascribed to

accumulation of uranium and the longer residence time of water here (Chung & Swart, 1990).

A secondary cause related to the lower pH in the phreatic zone which lower the concentration

of CO32- ions and hence raised the UO2(CO3)22-/CO32- ratio of the fluid.

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Boron and Sulphur Isotopes

The impact of marine and freshwater diagenesis upon the 34S (CAS) and 11B values has not

been investigated thoroughly, although dissolution and precipitation of carbonates within

regimes which are not anoxic does not appear to alter the 34S of the CAS (Gill et al., 2008).

There is a suggestion that the 11B values are altered in the freshwater phreatic zone associated

with the lower pH in this region (Stewart et al., Submitted).

Calcretes

Calcretes or caliche surfaces form in a variety of ways and represent the sub-aerial weathering

of carbonate surfaces (Esteban & Klappa, 1983; James, 1972; Rossinsky & Swart, 1993;

Rossinsky et al., 1992; Sheldon & Tabor, 2009; Watts, 1980) correlating with periods of

emergence (Beach & Ginsburg, 1982; Kievman, 1998; McNeill et al., 1988). Calcretes can form

on the surfaces of bare carbonate rocks and within soil zones (Fig. 10A and B). On bare

carbonate rocks, laminae form in response to dissolution of the surface by corrosive rain and

then the subsequent precipitation of carbonate minerals during drying events. During the

wetting events small depressions in the rock are colonized by microalgae and cyanobacteria

which photosynthesize and then decay, imparting an overall negative 13C signature to the

water. Evaporation causes the water to become isotopically enriched in 18O and therefore

these laminated carbonates and/or calcretes have negative 13C and occasionally positive 18O

values. Caliches found within the soil zone are characterized by root structures (rhizoliths)

(Klappa, 1980; Fig. 10C) and also have more negative 13C values, but are rarely enriched in 18O

(Rossinsky et al., 1992); this is because meteoric waters rapidly penetrate the soil layer thus

protecting the fluids from evaporation. Soil carbonates therefore tend to reflect patterns of

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global precipitation (Rossinsky & Swart, 1993; Sheldon & Tabor, 2009) as defined by previous

workers (Rozanski et al., 1993). The magnitude of 13C depletion within the soil zone has been

related to the nature of the vegetation and atmospheric pCO2. For example, the 13C values of

diagenetic carbonates from soil profiles in East Africa allowed the timing of grassland

development in the area to be determined as vegetation changes from C3 to C4 plants (Cerling,

1989; Cerling et al., 1977).

Palaeo-atmospheric pCO2 levels have been inferred from the 13C value of soil carbonates

based on the work of Cerling (1991) who recognised that soil CO2 is a mixture of atmospheric

and soil-respired components. This approach has been used by a number of workers to

estimate atmospheric pCO2 through time (Ekart et al., 1999; Montanez et al., 2007), although

the model is based on many different assumptions and the constancy of such assumptions

through time is difficult to verify. One of these assumptions is that the soil carbonate formed at

temperatures representative of the growing season. Although temperatures are recorded in

the 18O values (Dworkin et al., 2005), interpretation is complicated as a result of changing

18Ow values and the fact that carbonate seems to form mainly during dry periods (Breecker et

al., 2009). The solution to this problem might be to apply a technique such as the clumped

isotope thermometer (Passey et al., 2010) although some studies, supported by petrographic

analyses, have suggested that subsequent alteration of the soil carbonates may be a significant

problem (Budd et al., 2002) for their use as palaeoenvironmental indicators.

Calcretes and caliche crusts are typically elevated in Fe, Mn, Si and Al (Machusak & Kump, 1997;

Rossinsky et al., 1992). In the Bahamas such elements are derived from atmospheric dust

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originating in the Sahel region of Africa (Muhs et al., 2007; Prospero, 1990; Prospero et al.,

1970). Carbonates forming deeper in soil horizons (termed penetrative calcretes) also have

elevated Fe and Mn concentrations, but typically lower than calcretes forming near the surface

(Rossinsky et al., 1992). These elements are present not only substituting for Ca in the calcite

structure, but also as contaminant minerals (clays and oxides).

Vadose Zone

The vadose zone is defined as the portion of the rock or regolith above the water table.

Technically this zone also includes the calcrete and caliche portion of the soil profile and

consequently receives abundant contributions from the decay of OM and CO2 from the

respiration of roots. Rainwater entering the vadose zone contains only small quantities of

dissolved salts, derived from aerosols and atmospheric CO2. Here the water initially dissolves

local calcium carbonate until saturation is attained with respect to the ambient carbonate

mineralogy. In geologically young terrains, mainly composed of aragonite and HMC,

precipitation of LMC can occur before dissolution of aragonite and HMC is complete (Budd,

1988). Within the vadose zone, LMC precipitates as meniscus and pendant cements as the

water drips between grains (Halley & Harris, 1979), with the shape of the cements reflecting

the concave nature of the water surfaces (Fig. 10D). Allan & Matthews (1982) described this

zone as showing a wide range in 13C values, but with a rather constant and negative 18O value

(Fig. 11). Changes also occur in the composition of trace and minor elements during diagenesis.

Such changes usually involve the dissolution of metastable carbonates (aragonite and HMC)

releasing trace elements into the pore water, followed by precipitation of more stable phases

such as LMC. Because the vadose zone is located immediately below the soil horizon, cements

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forming in this region frequently have high concentrations of elements such as Mn and Fe (10

to 1000 ppm; Kievman, 1998). These elements can be derived from atmospheric dust in

locations remote from terrestrial influences (Rossinsky et al., 1992; Swart et al., 2014).

Fresh-water Phreatic Zone

The fresh-water phreatic zone is located below the water table and is characterized by both

lateral and vertical flow. Here isopachous equant cements grow into pore space filled with

water, forming a characteristic dog-toothed spar (Halley & Harris, 1979; Fig. 10E). At the water

table itself diagenetic reactions are particularly aggressive as OM, which has filtered through

the vadose zone, accumulates at the water-air interface. Here the concentration of oxygen is

lowered as OM is oxidized and sulphate reduction (utilizing the small amount of sulphate in the

meteoric fluid derived from aerosols) causes pH to decrease, and dissolution and precipitation

reactions to take place (McClain et al., 1992). Based on changes in the water chemistry (SO42-,

alkalinity, Ca2+ and Sr2+) the reaction zone in the study of McClain et al. (1992) appeared to be

localized to within 10 to 20 cm of the water table. The sediment in the majority of the

freshwater phreatic zone in that study was uncemented leading McClain to argue that a

changing sea level caused the active diagenetic zone to move through the sedimentary column

promoting cementation (Budd & Land, 1990; McClain et al., 1994). In contrast Budd et al.

(1988) suggested that the original aragonite and high-Mg calcite sediments within the

freshwater phreatic zone would eventually stabilize to LMC as a result of the mineralogical

drive (James & Choquette, 1982), the difference in solubility between metastable minerals such

as aragonite and HMC and LMC. Regardless of the mechanism of stabilization, within the fresh-

water phreatic zone the 18O values are still more negative (similar to the vadose zone; Allan &

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Matthews, 1982; Land, 1970). The 13C values are less variable and slightly more positive

compared to the vadose zone, although still isotopically negative relative to the original marine

sediments. Although it is generally believed that the freshwater phreatic zone is characterized

by the presence of cements as shown in Fig. 10E, muddier portions of the section frequently do

not show such cements, yet still possess relatively negative 13C and 18O values (Fig. 10F).

The Inverted J Pattern

The pattern of alteration of marine sediments by meteoric fluids has been described by

Lohmann (1987) as following an inverted ‘J’ pattern (Fig. 11) in 13C and 18O space. Original

carbonate sediments, which fall within a range of 13C and 18O values depending on their

origin (Fig. 5), undergo dissolution and precipitation reactions mediated by meteoric fluids

changing the bulk 18O of the rock towards more negative values. In an open system there is

abundant fluid moving through the rock and therefore the 18O value of the original carbonate

becomes quickly masked by the local meteoric fluids as the rock is altered. At the same time

the 18O value is altered, the 13C value of the diagenetic carbonate is influenced by oxidized

OM. As a result of the large amount of carbon in the sediment/rock compared to the water

(contrasting with oxygen), alteration of the 13C value takes significantly longer than the 18O

value. The evolution of the carbonate as traced on a 13C versus 18O plot shows the 18O value

moving into equilibrium with the groundwater at a specific temperature while the 13C values

span a wide range. This is the so-called inverted ‘J’ trend (Fig. 11). The vertical lines on this

diagram are described as the Meteoric Calcite Lines (MCL) and represent the 18O values of

calcite in equilibrium at a specific temperature with ground waters of different 18O values. As

the MCL varies as a function of the 18O content of the rainfall, it also varies relative to latitude

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and ambient temperature (Hays & Grossman, 1991). For example, in the more recent record,

the 18O value of calcretes varies throughout the Caribbean (Rossinsky & Swart, 1993) relative

to the amount of rainfall and its 18O value. As a final observation many diagenetic pathways

might not follow the inverted ‘J’ if the original sediment has a 18O value close to or even more

negative than that represented by the MCL. In addition, the transition from the original

sediment might take place rapidly and might not be represented in the rock record. This is

actually the case in the data shown in Fig. 11.

Mixing-zone

The mixing-zone describes the transition between fresh and marine waters occurring along

coastlines and small oceanic islands. As a result of the non-linear behaviour of activity

coefficients of ions in fluids of varying salinity, the mixing of fresh and saline waters, which are

both oversaturated with respect to carbonate minerals such as calcite, can produce waters

which are undersaturated relative to the same minerals (Badiozamani, 1973). Such behaviour

can cause extensive dissolution (Back et al., 1986; Whitaker & Smart, 1997). Mixing-zone

associated caves are therefore common within carbonate terrains that are close to sea level.

These mixing-zones might also be regions of dolomite formation because while fluids in the

mixing-zone are undersaturated with respect to calcite, these same fluids are still highly

supersaturated relative to dolomite (Badiozamani, 1973). While this notion was briefly very

popular with several authors claiming to have documented the occurrence of mixing-zone

dolomites based on stable C and O isotopes and minor elements (Humphrey, 2000; Ward &

Halley, 1985), the evidence is not compelling. The 18O values for example cannot be used to

elucidate the process of dolomitisation because of uncertainties (discussed previously)

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regarding the most appropriate equilibrium relationship to use. The mixing-zone mechanism of

dolomitisation has fallen out of favour principally as a result of the absence of dolomite in most

modern day mixing-zone environments (Budd, 1997; Hardie, 1987; Machel & Mountjoy, 1986;

Melim et al., 2004). Notwithstanding the arguments regarding its significance, Allan &

Matthews (1982) have shown the mixing-zone to be characterized by a strong linear covariation

between 13C and 18O values. It has been suggested that this results from alteration in fluids

with progressively more positive 18O values as one progresses from fresh water to marine

water through the mixing zone.

Blue Holes

Blue Holes are collapse features observed in modern carbonate terrains such as the Bahamas;

they are believed to form as a result of dissolution related to the movement of the mixing-zone

through the carbonate platform as sea-level oscillates. Modern Blue Holes in the Bahamas have

a layer of freshwater water overlying seawater. Unlike groundwater settings, the OM in the

Blue Holes appears to sink through the freshwater, accumulating at the seawater interface

(Jones et al., 2014). Here the OM is rapidly degraded utilizing all the available oxygen and then,

with help from the sulphate in the underlying seawater, copious amounts of H2S are produced.

The presence of H2S, combined with the naturally corrosive characteristics of the mixing zone,

produces enhanced dissolution in these caves as the interface oscillates under the influence of

tides and natural sea-level changes.

Other Mixing Phenomenon

In the previous example, the mixing zone has been defined to represent interacting solutions

with varying salinities. However, non-conservative behaviour can also take place when two

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fluids, both saturated with respect to a carbonate mineral such as LMC and with similar ionic

strengths, but with different CO2 fugacities, are mixed. Undersaturation with respect to

carbonate minerals occurs as the changes in the concentrations of the various carbonate ions

are not conservative. Such behaviour is common in groundwaters charged with CO2 after

passing through soil zones and then mixed with a second body of water with a lower pCO2.

Evaporative Waters

Climate, principally the amount of rainfall, can also control surface related diagenesis. Two

good modern examples are the Bahamas and the Persian Gulf, both located at about the same

latitude. In the Persian Gulf, rainfall is very low and evaporation extremely high. This leads to

the development of hypersaline conditions in tidal flats and lagoons and the formation of a

suite of minerals associated with the evaporation of seawater (evaporites). In contrast, the

Bahamas experiences greater precipitation and low evaporation. Consequently, evaporites are

rarely found associated with modern sediments in the Bahamas. Despite the prevalence of

evaporites in arid settings, saline fluids can also form in water bodies with restricted circulation.

Although such fluids might reach more modest salinities (40 to 120), they are highly saturated

relative to carbonate minerals and in some instances are associated with higher amounts of

organic degradation or higher rates of photosynthesis. Such saline waters could be dense

enough to sink or reflux through underlying sediments causing diagenetic alteration as they

migrate. Generally such fluids would have positive 18O values, but might have either positive

or negative 13C values depending upon the relative amounts of photosynthesis and

degradation of organic material (Swart et al., 2009).

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During evaporation of seawater calcium carbonate is continually precipitated with the removal

of Ca2+ and CO3- from solution, the remaining fluids are enriched in Mg2+, Na+, Cl-, K+ and SO42-.

During the later stages, minerals such as gypsum, dolomite and halite can eventually be formed

along with aragonite and a host of other evaporative minerals (Schreiber & El Tabakh, 2000).

Because the distribution coefficient for the incorporation of Sr into evaporite minerals is lower

than one (Holser, 1979), the amount of Sr in the residual brine increases during precipitation of

these minerals and celestite (SrSO4) can be found as an accessory mineral. Although other

trace elements can also change concentration depending upon their distribution coefficients

and the minerals being formed, there are no data concerning the values of the distribution

coefficients for elements such as Mn, Fe and Mg into common evaporite minerals.

Evaporite associated carbonates are frequently dolomitised as a result of the high Mg/Ca ratio

of the evaporative fluids (McKenzie, 1981; McKenzie et al., 1980), the presence of abundant

OM which promotes high alkalinity (Schreiber et al., 2001) and certain types of bacteria. All of

these conditions are known to promote the formation of dolomite and protodolomite

(Vasconcelos & McKenzie, 1997; Warthmann et al., 2000). Since depletion of O2 and the use of

sulphate as an oxidant of OM leads to the production of CO2 with negative 13C values,

dolomites (and other carbonates) formed under such conditions tend to have more negative

13C values than marine carbonates. The 34S values of evaporate associated sulphate minerals

are generally believed to reflect the 34S of the sulphate in the oceans. However, wide ranges

in 34S values found in evaporite deposits during certain times probably reflect processes of

sulphate reduction and the oxidation of OM.

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Marine Diagenesis

Sea Floor

Evidence of the influence of early marine diagenesis on the 18O and 13C values of carbonate

particles is poorly documented in the literature. Processes such as micritisation and the

precipitation of aragonite and HMC cement undoubtedly affect the 18O and 13C values of the

original sediments, but such processes are difficult to assess as the cements and areas of

alteration are generally too small to allow effective sampling without cross-contamination of

the original material. It is likely, however, that the addition of cements to a marine framework

originally possessing more negative 13C values will tend to enrich the samples in 13C (as

inorganic aragonite and HMC cements tend to have more positive 13C values; Aissaoui, 1988;

Grammer et al., 1993). Studies that have sampled such cements confirm their more positive

13C values (compared with normal skeletal aragonite) and 18O values close to equilibrium

(Braithwaite & Camoin, 2011). Processes influenced by photosynthesis, such as those adding

cement within stromatolites, might have more positive values as a result of the preferential

removal of 12C or alternatively more negative values as a result of CO2 produced during

sulphate reduction (Andres et al., 2006). The 18O value of the inorganic cements is also likely

to be different depending on the nature of the original grains and the temperature of

precipitation. The trace and minor element concentrations might also be altered by

micritisation and precipitation of aragonite cements. Micritised borings are often filled with

HMC and therefore would be lower in Sr and higher in Mg, if the grain was originally aragonite.

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The borings might also have a different 18O and 13C value depending on the nature of the

filling (aragonite or HMC) and the diagenetic processes involved.

During periods of non-deposition, cementation takes place near or at the sea floor resulting in

the formation of hardgrounds and firmgrounds. In cases where carbonate cementation occurs

in colder and deeper waters than the environment in which the sediment formed, the 18O

value of the subsequent cements will be more positive than the host sediment, overprinting or

modifying the original bulk sedimentary signature. The 13C value of the precipitate may also be

positive as inorganic aragonite precipitated in equilibrium with normal ocean water would be

expected to have a 13C value of +4‰. This process has been argued to account for positive

13C and 18O values at hardground surfaces in ancient rocks (Marshall & Ashton, 1980).

However, cementation in hardgrounds can also lead to more negative 13C values as products

of microbial decomposition become incorporated into the diagenetic cements. An example of

this can be found in cores drilled in the Bahamas (Fig. 12), where a surface at a depth of 536 m

in the Clino core, representing a hiatus of 2 to 3 Myr, became cemented with carbonates

possessing more negative 13C and more positive 18O values than those of the original

sediment (Swart & Melim, 2000; Fig. 13). These sediments were originally derived from the

surface of GBB and had 13C and 18O values in equilibrium with shallow, warm water with

relatively positive 13C values. During the period represented by the hiatus, they became

burrowed and enriched in OM. This OM was eventually degraded by oxic and anoxic processes

leading to the production of CO2 depleted in 13C and cementation in cooler waters. Additional

examples of negative 13C values in hardgrounds are provided by Dickson et al. (2008). In this

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case the authors postulated that the more negative 13C values were the product of sulphate

reduction similar to the model proposed by Swart & Melim (2000).

Geochemical Pore water Profiles in Marine Sediments and Diagenetic

Implications

The Deep Sea Drilling Project (DSDP), the Ocean Drilling Program (ODP) and their successors

have substantially increased current knowledge of the diagenesis of carbonates in the deep

marine environment. Pore waters can be retrieved from such sediments by squeezing the

unconsolidated materials in a press or through various types of water sampling devices as

described by Martin & Sayles (2003). The analysis of the geochemistry of such pore waters is

potentially much more powerful than examining the sediments themselves. For example, if

dolomite is formed in a sediment with 50% porosity utilizing the local CaCO3 and Mg2+ derived

from the pore water, then only ca 0.1 % of dolomite could be generated utilizing all the

available Mg2+ in the pore fluid. Such a drastic change in the pore water chemistry could easily

be measured, but it might be very difficult to find the corresponding 0.1% of dolomite in the

bulk sediment.

Marine Burial

In the upper portion of the sedimentary column, close to the sediment/water interface,

oxidation of OM proceeds using O2 as the terminal electron acceptor. This process generates

CO2 which, in turn, causes the pH to decrease and the pore waters to become undersaturated

with respect to the more soluble carbonate minerals. During low rates of sedimentation the

products of oxidation escape into the overlying water column and the pore fluids become

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resupplied by new O2 as the OM is consumed. The net effect is a decrease in the concentration

of OM, but little physical or chemical alteration of the sediments, although some sea-floor

cementation may take place under low rates of sedimentation. As rate of sedimentation

increases, the supply of O2 and the removal of CO2 are limited by diffusion. The most

favourable alternate oxidant under anoxic conditions is Mn4+, followed by NO3- and Fe3+.

Compared to O2, which provides energy of 3190 kJ for every mole of glucose oxidized, these

alternate oxidants yield ca 3000, 2750 and 1400 kJ mol-1, respectively (Froelich et al., 1979).

However, in practice the concentrations of these species (Mn4+, NO3- and Fe3+) in the modern

oceans are very low and they are not normally important in the degradation of OM.

Volumetrically the most important alternative oxidant to O2 is SO42-, even though the energy

yield of 380 kJ mol-1 is considerably lower than that of the other oxidants. Typically in modern

sediments there is a transition from the zone in which oxygen is the principal oxidant to one

where sulphate is utilized by bacterial sulphate reduction (BSR) to degrade OM [anoxic sulphic

zone (Berner, 1981)]. Regardless of the oxidant utilized, all these processes contribute CO2 with

negative 13C values, similar to the parent OM. In the SO42- reduction zone abundant H2S is also

produced. Once SO42- has been completely utilized the oxidation of OM proceeds through the

process of methanogenesis [anoxic non-sulphic zone (Berner, 1981)] which results in a large

fractionation between CH4 and CO2, resulting in CO2 with positive 13C values (+10‰) and CH4

with very negative 13C values (-40 to -60‰; Games et al., 1978). It is believed that

methanogenesis terminates once the geothermal gradient elevates the temperature above the

viable limit for BSR (Coleman et al., 1979). Below this depth the OM degrades under the

influence of thermal processes which produce isotopically negative CO2 and CH4, as well as

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higher molecular weight alkanes (CNH2N+2). Methane and other thermogenic gases can migrate

upwards through the sediments and be utilized by bacteria within the BSR and oxic zones

(methanotrophy) and in certain conditions form gas hydrates.

The transitions as described above occur not only within the pore water in sediments, but also

in stratified anoxic water bodies such as lakes or basins, for example, the Black Sea (Murray et

al., 1991) or Cariaco Basin (Scranton et al., 1987). During certain periods of Earth history, such

as the Cretaceous, the number of anoxic basins was significantly greater than today (Jenkyns,

2010) and during the early Earth all oceans were anoxic (Busigny et al., 2014; Lyons et al.,

2014).

Limitations imposed by the supply of oxidants have important consequences regarding the

extent to which carbonate sediments can be influenced by the degradation of OM. Consider 1 g

of carbonate sediment composed of aragonite (density = 2.9 g/cm3) with 50% porosity. Based

on a seawater SO42- concentration of 28 mM, within the pore space there would be

approximately 9.6*10-7 moles of SO42- available for the oxidation of any OM present in the pore

fluids. That would produce twice that amount of HCO3-. In contrast there are 0.01 moles of C in

the 1 g of carbonate sediment. By simple mass balance calculation it is apparent that the

oxidation of OM within carbonate sediment cannot alter the 13C value of the original sediment

significantly unless additional oxidants are supplied. Although such a supply is possible through

the process of diffusion, considerable time would be necessary to allow sufficient SO42- to

diffuse downward from the overlying seawater. In addition, the greater the depth over which

diffusion takes place, the less likely it will be that the 13C value of the bulk sediments will be

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impacted by the oxidation of OM. This implies that in most sediments deposited in the marine

realm, it will be difficult to significantly alter the 13C value of the original sediments, even if

100% of the original sediment is dissolved and reprecipitated. The exception to this would be

sediments close to non-depositional surfaces, particularly if such surfaces represented a long

hiatus, or in sediments with a low rate of deposition. Alternatively if carbonates are a minor

contributor to the sediment, then the 13C value of any diagenetic component will be strongly

influenced by the oxidation of OM. If the system were open, sufficient electron acceptors may

be available to oxidize all the OM. However, by definition, in an open system the products of

diagenesis would be removed from the source of oxidation. The influence on the diagenetic

carbonate would be balanced by the rate of alteration and the rate of fluid flow.

In the majority of cases the geochemical parameters in the pore waters of fine-grained and

muddy sediments are controlled by diffusive rather than advective processes. A diffusive

profile is one in which the transport of an element is governed by a difference in concentration

between the interstitial pore waters and an external source such as the overlying seawater or

an underlying reservoir. An ion diffuses from a region of high concentration to one of low

concentration as a result of the concentration gradient modified by diagenetic reactions. The

distribution of a species can be described by the general diagenetic equations as proposed by

Berner (1980). In this equation the change of the concentration of any component (C) as a

function of depth (z) and time (t), bulk diffusion coefficient (Db) and diagenetic change (J) is

expressed by Eq. 2:

C   C  (vC)
  Db z   z  ( J i ...J n ) (2)
t z  

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The dominant role of diffusion can be recognised not only in the shape of the profiles of minor

and trace elements in pore waters squeezed from sediments, but also in the geothermal

gradients measured in the sediments. There are instances, however, where advective

processes control the temperature and geochemical profiles. An advective process is one in

which there is movement of fluids driven by external processes unrelated to the actual

concentrations of ions. While a certain amount of advection occurs during compaction, this is

slow relative to diffusion. Significant advective processes can be recognised as either concave

upward profiles of temperature or elemental composition, suggesting movement of fluid out of

the sediments, or concave downward profiles, indicating movement of fluid into the sediments.

Typically in oceanic sediments, these processes occur associated with hydrothermal processes

near spreading centres (Kastner et al., 1986), in carbonate platforms, where large-scale

circulation is perhaps driven by differences in temperature between the platforms and the

adjacent oceans (Kohout, 1967), density difference between the waters on the platform and

the adjacent oceans (Simms, 1984) and along continental margins, where there is a hydrological

head associated with the adjacent mainland (Sayles & Manheim, 1975). The movement of

seawater throughout carbonate platforms by Kohout convection and reflux of brines with

slightly higher salinity than normal seawater has been discussed in numerous, mainly

theoretical studies (Jones et al., 2002; Simms, 1984). Such processes are potentially more

important in controlling diagenesis than diffusive ones, in that advection is able to supply

reactants needed for the alteration of sediments as well removing the products of these

reactions at significantly higher rates.

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Diagenetic processes taking place within buried sediments can be identified using ratios of

diagenetically active elements relative to a conservative element such as Cl- or ratios of

diagenetically active elements such as Ca2+, Sr2+ and Mg2+ to one another, i.e. Mg2+/Ca2+ and

Sr2+/Ca2+. The excess or deficit of an element or species (x) relative to a conservative element in

seawater such as Cl- can be calculated using Eq. 3:

(3)

If the concentration of an element behaves in a conservative manner then the excess is zero; if

the element is consumed then the excess is negative and if the element is generated the value

is positive. An example of this approach using data from ODP Site 1005 located on the margin

of GBB (see Fig. 12 for location) is shown in Fig. 14. At this site, excess (deficit) concentrations

of all minor species cluster near zero in the upper 50 m of the core. This indicates either an

absence of geochemical reactions in this portion of the core or a rapid flushing of fluids so that

the products of diagenesis are removed. Below this depth Ca2+, Mg2+ and SO42- all decrease in

concentration, while the alkalinity increases. Decreases in Ca2+ and Mg2+ reflect the

precipitation of LMC and dolomite, while changes in alkalinity and sulphate concentration

indicate the oxidation of OM by sulphate:

(4)

In the absence of: (i) reactions which utilize HCO3-; and (ii) preferential diffusion of HCO3-

relative to SO42-, the deficit of SO42- should be 28 mM and the excess of HCO3- 56mM (utilizing

Eq. 4; Fig. 15). The precipitation of carbonate minerals can also be seen in changes in the ratios

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of Mg2+/Cl-, Ca2+/Cl-, Mg2+/Ca2+and Sr2+/Ca2+ in the pore fluids (Figs 16 and 17). In these figures,

data are plotted from two localities, one from cores in Florida Bay (Fig. 12; Burns & Swart,

1992; Swart et al., 1987a) and from ODP Site 1005 (Kramer et al., 2000) drilled at a water depth

of 888 m. Despite the difference in core locations, the diagenetic reactions which occur at

these two localities are similar. For example, the precipitation of LMC at both sites results in a

reduction of the Ca2+/Cl- ratio (as Ca2+ is consumed) and an increase in the Sr2+/Ca2+ ratio (as

Sr2+ is preferentially rejected as a consequence of the low distribution coefficient for the

formation of LMC; Fig. 16). Changes in the concentration of Mg2+ and Ca2+ resulting from the

precipitation of dolomite also take place, but will depend upon the stoichiometry of the

reaction involved. In the case of dolomitisation using Eq. 5 there will be a reduction in the

Mg2+/Cl- and Ca2+/Cl – ratios and an increase in the Sr2+/Ca2+. The formation of dolomite by this

equation should not change the Mg2+/Ca2+ ratio of the pore fluids as Mg2+ and Ca2+ are used in

equal molar amounts (Fig. 16):

(5)

In contrast dolomitisation utilizing Eq. 6 involves the dissolution of existing calcium carbonate

and therefore this reaction increases the Ca2+/Cl- ratio while decreasing the Mg2+/Ca2+ ratio:

(6)

Pathways for other possible diagenetic reactions such as aragonite and HMC dissolution and

LMC and dolomite precipitation are shown in Figs 16 and 17.

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Within sediments the formation of minerals such as calcite and dolomite causes Sr to be

preferentially excluded and consequently this ion builds up in the pore fluids. An increasing

Sr2+/Ca2+ ratio is therefore a consequence of the formation of either, or a combination, of these

minerals. Changes in the Mg2+/Cl-, Ca2+/Cl- and Mg2+/Ca2+ accompanying an increase in

Sr2+/Ca2+ can suggest which of the reactions are actually taking place. Such changes are readily

visible by using plots of the Mg2+/Cl- with respect to Ca2+/Cl -, SO42-/Cl- and other ions as

outlined by Swart & Kramer (1998).

Geochemical Pore water Profiles in Marine Sediments and Diagenetic Implications

As a first approximation, the diagenesis of carbonate buried in the marine realm can be

separated into those which overly basaltic basement, those found along the margins of

carbonate platforms and those located along the margins of continents.

Deep Oceanic Carbonates

Modern oceanic sediments are dominated by LMC producing organisms. At the seawater

interface the sediments form an ooze and are uncemented with porosities of 60 to 70% being

typical. With increasing depth the sediments become compacted, thereby reducing porosity

and the processes of dissolution and precipitation eventually transform the ooze into a hard

rock.

Calcium and Magnesium

Oceanic carbonates typically show an increase in the concentration of Ca2+ and a decrease in

Mg2+ in the interstitial pore fluids with increasing depth [between 1 m and 300 m below sea

floor (mbsf)]. These changes arise partially as a result of dissolution and precipitation within

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carbonate sediments, including dolomitisation, but principally as a result of alteration of basalts

at the interface between the sediments and the underlying igneous rocks (McDuff & Gieskes,

1976). In addition secular changes in the Ca2+ and/or Mg2+ of seawater can modify the profile

(Higgins & Schrag, 2012). Reactions at the basalt-sediment interface produce Ca2+ and consume

Mg2+ setting up a strong diffusion gradient in the interstitial pore water from the sea floor to

the sediment-basalt boundary (McDuff, 1976; McDuff, 1981), thicknesses of hundreds of

metres. Such reactions preferentially incorporate 26Mg leading to a decrease in the26Mg

values of the pore waters. The two examples shown in Fig. 18 are taken from DSDP Site 504

(Mottl et al., 1983), situated above basaltic basement on the Costa Rica Rise and DSDP Site 541

(Gieskes et al., 1984), located on the Barbados accretionary prism. Interstitial-pore fluid

concentrations of Ca2+ at both sites increase from seawater values of ca 10mM to ca 50 to 60

mM, while Mg2+ values decrease from 55 mM to less than 20 mM close to the boundary

between the sediment and the underlying basalts. The 44Ca in pore waters in oceanic

sediment typically become more negative as the carbonates are altered to LMC and impart

their original negative 44Ca value to the pore fluids (Fantle & DePaolo, 2007) .

Strontium

Although Cenozoic oceanic sediments mainly consist of foraminifera and coccoliths composed

of LMC, such biogenic carbonates tend to be metastable compared to inorganic LMC and over

time dissolve and precipitate forming inorganic LMC (Baker et al., 1982). Oxidation of OM by

sulphate causes initial undersaturation (Ben-Yaakov, 1973) promoting dissolution and

consequent precipitation of the more stable inorganic carbonate phases. With depth , there is a

steady increase in the concentration of Sr2+ in the pore fluids (Baker et al., 1982; Gieskes, 1976)

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(Fig. 18), superimposed on this are increases in Ca2+ and decreases in Mg2+ as described earlier.

The Sr2+/Ca2+ ratio in the pore fluids can either decrease, increase or remain the same,

depending on the magnitude of Ca2+ and Sr2+ changes. The concentration of Sr2+ in the pore

fluid increases until it reaches saturation with respect to the mineral celestite (SrSO4).

Consequently celestite is frequently found coincident with the Sr2+ maximum in the pore water

(Baker & Bloomer, 1987), suggesting direct precipitation from the pore fluids. In the absence of

sulphate reduction, the maximum concentration of Sr2+ which can be attained in the pore fluid

is ca 700 M (normal seawater Sr2+ concentration ca 90 M). In instances where there is

abundant sulphate reduction the concentration of Sr2+ can reach much higher values (see

below). In some instances external sources of sulphate diffuse into the reduction zone from

underlying sediments (Kramer et al., 2000). In these cases Sr2+ in the initial pore fluids is

removed (as a result of the precipitation of celestite) causing a sink of Sr2+ deeper in the

sedimentary column. Strontium in such situations can diffuse both upward and downward

from the Sr2+ maximum zone (Gieskes et al., 1986). The 87Sr/86Sr ratio is also an indicator of Sr2+

diffusion within the interstitial pore fluids. Strontium released to the interstitial pore fluids

initially has a 87Sr/86Sr ratio similar to that of the local sediment. As the Sr2+ diffuses up or down

through the sediment column this signature changes along with the transported Sr2+, not only

because the 87Sr/86Sr ratio is different, but also as a result of preferential diffusion of 87Sr

relative to 86Sr. Hence if one considers the last ca 60 Myr, during which the 87Sr/86Sr ratio of the

oceans has become increasingly more radiogenic towards the present day, then in the

sediments above the pore fluid Sr maximum, the 87Sr/86Sr ratio of the pore fluids is typically less

radiogenic than the sediments themselves. Below the region of maximum concentration the

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87
Sr/86Sr ratio is more radiogenic (Elderfield & Gieskes, 1982; Gieskes et al., 1986). This effect

can be seen in the example from Site 541 (Fig. 18). Here the concentration of Sr2+ reaches a

maximum of 400 M around 300 mbsf. This corresponds to the point at which Sr2+ diffuses

upwards imparting a less radiogenic (older) signal to the pore waters in the younger sediments

and a more radiogenic signal to the sediments below.

Oxygen Isotopes

The18Ow values in pore fluids decrease with increasing depth at oceanic sites (Fig. 18), as a

result of the alteration of the underlying basalts to clay minerals (Lawrence et al., 1975). This

trend for the pore fluids to assume more negative 18Ow values is, however, ameliorated to

some extent by the tendency for pore waters to become enriched in 18O as a result of

dissolution and precipitation of carbonates at higher temperatures (Killingley, 1983).

Carbonate precipitated along such a gradient would be depleted in 18O, not only as a result of

changes in the 18Ow values, but also as a result of the increasing temperature. In the upper

portion of oceanic sediments there is a slight increase in the 18Ow values which is not related

to carbonate dissolution and reprecipitation reactions. This change is a result of the increase in

the 18Ow value of glacial seawater which is now preserved in the pore fluid profile. With time

this increase will diffuse away, but the current 18Ow profile has been used to model the change

in the 18Ow values of the oceans during the last glacial period (Adkins & Schrag, 2003; Schrag,

1996).

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Organic Matter Oxidation

The extent to which SO42- is utilized in the pore fluids depends upon the supply of OM to the

location and the rate of sedimentation. In the open ocean where sedimentation is low, OM is

effectively oxidized near the sediment/water interface and the proportion of OM which

survives is reduced. As a consequence, decreases in SO42- and increases in alkalinity in the pore

waters are low. For example, at Site 709 drilled in the Indian Ocean, SO42- decreases only

slightly in the upper portions of the sedimentary column and the concentration of SO42- and

alkalinity remains close to seawater values throughout a sedimentary sequence of 250 m

(Backman et al., 1988). As a result, rates of dissolution and precipitation are relatively low

leading to only a small increase in the concentration of Sr2+ in the pore fluids. As this particular

site overlies basaltic basement there is still a large increase in the concentration of Ca2+ and

decrease in Mg2+ driven by reactions at the sediment-basalt interface.

Implications for Diagenetic Carbonates

Diagenetic carbonates formed during the burial of deep ocean carbonates tend to show

decreasing 18O values with burial, both as a result of the increased temperature and as a

consequence of the decreasing 18Ow values of the pore fluids caused by basalt–water

interaction. In contrast, the 13C values of the diagenetic carbonate should not be significantly

different than the 13C values of the original sediments (see previous discussion). Pore water

Sr2+ concentrations increase with depth in oceanic carbonates, but a combination of more

rapidly rising Ca2+ concentrations and an increasing temperature would tend to lower the Sr

concentration of the diagenetic calcite. The combination of a low 18O value and low Sr

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concentrations in a calcite might be incorrectly interpreted as indicating that such sediments

were affected by meteoric diagenesis.

Using Pore water profiles for Estimating the Amount of Dissolution and Precipitation

Various workers have utilized pore water profiles, particularly Sr and 18Ow but more recently of

Ca, Mg and U isotopes, in order to understand and model rates of recrystallization in deep sea

carbonates (Baker et al., 1982; Fantle et al., 2010; Killingley, 1983; Lawrence et al., 1975;

Richter & DePaolo, 1987; Schrag et al., 1995; Stout, 1985). The most recent review of the

subject provides an excellent overview (Fantle et al., 2010).

Strontium: The concentration of Sr2+ in the pore waters has been used to estimate the amount

of recrystallization in oceanic sediments, by comparing the total amount of Sr lost from the

sediments over a defined period of time with the total amount of Sr available from a given

amount of sediment (Baker et al., 1982). An alternative model is to compare the expected

Sr/Ca (or Mg/Ca) ratio of the sediments with the expected values based on the Sr, Mg and Ca

concentrations in the pore water. At the sediment-water interface the Sr/Ca in the sediment

will be substantially different to that predicted from the pore waters [using a DSr of 0.04 (Baker

et al., 1982)], but with depth these values will converge and complete replacement of the

sediments by a diagenetic derived calcite is considered to have occurred when these two values

become approximately equal. One-dimensional reactive transport models applied to Sr

enabled better constraints to be placed on the diagenesis and hence accurately predict

dissolution precipitation reactions and the evolution of 87Sr/86Sr (Richter & DePaolo, 1987).

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Oxygen Isotopes: The 18O value of the sediments and the pore waters can also be used to

estimate the amount of dissolution and reprecipitation experienced by carbonate sediments

(Lawrence et al., 1975; Schrag et al., 1995; Stout, 1985). In Eq. 7 the 18O value of the pore fluid

(δwe) after a specific amount of alteration can be calculated:

(7)

In this equation Mc = mole fraction of oxygen in the carbonate, Mw = mole fraction in the pore

fluid, C = 18O [SMOW (standard mean ocean water)] of the initial sediment, w = initial 18O

value of the water, R = amount of transformation from the parent mineralogy and =

fractionation factor between calcite and water at a specific temperature (O'Neil et al., 1969).

Using this approach and assuming that the 18O values of the present pore water profile has

been similar throughout the history of the sediment, rates of recrystallization, porosity and

geothermal gradients can be varied, and the predicted 18O value of the sediment matched to

that actually measured.

Calcium Isotopes: A more recent approach to estimating recrystallization has been through the

application of similar models as used by Richter & DePaolo (1987), but applied to the 44Ca

values of pore waters and sediments (Fantle & DePaolo, 2007). The method relies on the

difference between the equilibrium 44Ca values in Modern marine carbonates which show a

fractionation of 44Ca of 0.9987‰ compared with carbonates which are dissolved and

reprecipitated in the deep sea, a process in which there is no fractionation (= 1.000). Based

on data from ODP Site 807A from the Ontong Java Plateau, Fantle & DePaolo (2007)

determined that rates are very high in relatively young sediments (30 to 40% Myr-1), decreasing

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to lower rates (2 to 3% Myr-1) in sediments older than 2 Ma. Such rates are considerably higher

than those calculated using the methods of Baker et al. (1982) for DSDP Site 288 (Baker et al.,

1982) and ODP Sites 709 (Swart & Burns, 1990).

Slope Carbonates with Platform Derived Sediments

Carbonate sediments deposited on the margins of carbonate platforms, below the extent of

glacial-eustatic sea-level changes, differ in their diagenetic behaviour compared to oceanic

carbonates for several key reasons. First, the sediments contain mixtures of different forms of

calcium carbonate [aragonite, HMC and LMC; periplatform sediments (Schlager & James, 1978)]

and are therefore more reactive than pelagic LMC. Second, they frequently have higher

depositional rates and higher concentrations of OM (1 to 2 wt%; Crevello et al., 1981)

compared to oceanic settings. Third, they are situated significant distances above basement,

thus limiting the influence of basement-carbonate interactions upon the pore fluids and the

diagenetic minerals produced. Despite these differences, at least some of the geochemical

patterns observed in the pore fluids from oceanic carbonates are also observed in pore fluids

from carbonate platforms. As in the case of oceanic sediments, oxidation of OM by sulphate in

the pore waters causes initial undersaturation with respect to biogenic minerals. In contrast to

oceanic sediments periplatform sediments frequently have more than one carbonate phase and

therefore once the fluids are below the saturation state of aragonite and HMC, dissolution is

able to proceed until all of the metastable phases have been altered to LMC, even in the

absence of the continued oxidation of OM. This mechanism is similar to that described in

carbonate sediment exposed to freshwater saturated with respect to LMC. The dissolution and

reprecipitation of carbonates is readily visible in thin sections taken from sites such as Site

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1003, drilled during Leg 166 adjacent to GBB (Fig. 19). At this site sediments near the surface

(Fig. 19A) are composed of mixtures of pelagic and platform-derived material and are

unconsolidated. With increasing depth the sediments start to dissolve and cements appear

extending into the interparticle pore space and mouldic porosity (Fig. 19B). Early cemented

portions are resistant to compaction (Fig. 19C), but susceptible to further alteration and

fracturing (Fig. 19D). During burial dolomite can be formed either replacing original sediments

or precipitating in voids (Fig. 19E). During deep burial the sediments become compacted and

porosity is filled with cement (Fig. 19F).

Calcium and Magnesium

The concentrations of Ca2+ and Mg2+ in the interstitial fluids in platform derived sediments

behave differently when compared to oceanic sediments. This is because, in contrast to the

continued increase of Ca2+ and the decrease in Mg2+ in pore fluids of oceanic sediments,

periplatform sediments are influenced to a greater extent by local diagenetic reactions rather

than diffusion induced by reactions at the basalt–sediment interface. The concentrations of

Ca2+ can decrease in zones where there is abundant carbonate precipitation, while in other

zones the concentration increases as a result of dissolution. Concentrations of Mg2+ decrease

as a result of the precipitation of dolomite (Kramer et al., 2000). As 26Mg is discriminated

against during dolomitisation, the 26Mg of the pore waters frequently increase at sites of

dolomite formation (Higgins & Schrag, 2010). At other sites where dolomite is not forming, the

26Mg of the pore waters decreases with depth as a result of preferential incorporation of 26Mg

into a mineral phase suggested to be clay (Higgins & Schrag, 2010) .

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In the three examples from oceanic sites presented in this paper [Site 817 from Queensland

Plateau, Site 823 adjacent to the Great Barrier Reef (GBR) and Site 1005 adjacent to the GBB]

(Fig. 19A and B), the Ca2+ decreases initially reflecting a balance between precipitation and

dissolution and then gradually increases to values between 20 mM and 30 mM. Magnesium

also shows variable behaviour probably relating to the formation of dolomite.

Strontium

Similar to oceanic sediments, Sr2+ in the pore fluids of slope carbonates increases with depth.

However, because both aragonite and HMC initially have much higher concentrations of Sr than

coccoliths and foraminifera which dominate typical pelagic deposits, there is potentially much

more Sr available for release into the pore fluids. In addition, HMC and aragonite are meta-

stable minerals and hence are more soluble than LMC. Concentrations of Sr2+ in periplatform

sediments therefore tend to rise more quickly with burial than in oceanic sediments. Similar to

oceanic sediments, the maximum concentration of Sr2+ which can be reached in the pore fluids

is dictated by the ion activity product of celestite. This mineral is more prevalent in platform

derived sediments than oceanic sediments (Baker & Bloomer, 1987; Swart & Guzikowski, 1988).

Frequently sediments which have been partially cemented become fractured and celestite

precipitates along the fractures. In other instances the celestite forms cement which

encompasses altered and unaltered sediments. Similar mechanisms were probably also

responsible for the formation of celestite in ancient carbonates (Yan & Carlson, 2003). As rates

of deposition and concentrations of OM are higher in periplatform than oceanic sediments, the

concentration of SO42- in the pore waters can frequently become completely exhausted

allowing the concentrations of Sr2+ in the interstitial pore fluids to reach very high levels. Three

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examples are used to illustrate this behaviour (Figs 20 and 21). Ocean Drilling Program Site 817

is situated off of a submerged carbonate platform (Queensland Plateau) located to the east of

the GBR. This site shows low initial concentrations of aragonite and only small decreases in the

interstitial SO42- concentrations. Consequently the Sr2+ in the interstitial pore water reaches

values similar to those seen in oceanic sediments. As a result of the low concentration of Ca2+

in the pore waters, the predicted Sr of any diagenetic calcite formed from these pore fluids

reaches values in excess of 4000 ppm (Fig. 20A). At ODP Site 823 located adjacent to Site 817,

in the trough between the Queensland Plateau and the GBR (Fig. 20B), the concentration of

sulphate in the interstitial pore fluids falls to values close to zero. As a consequence the

concentration of Sr2+ reaches very high values (3 mM). The predicted Sr concentration of

diagenetic calcite is in excess of 8000 ppm. The final example, from the Bahamas (ODP Site

1005), is also one in which SO42- decreases to zero throughout a large portion of the section

(Fig. 21A). The high amount of initial aragonite in the section and the processes of carbonate

dissolution and precipitation allow the concentration of Sr to reach up to 5 mM. The predicted

concentration of Sr in any diagenetic carbonate formed from these pore fluids would be 4% and

these concentrations are reflected in diagenetic cements at ODP sites such as Site 823 (Dix,

1995)

Although the behaviour of Sr-isotopes in the pore fluids from platform derived sediments

appear to be similar to that seen in the ocean carbonates, there is a difference which relates to

the rapid rates of dissolution and precipitation exhibited by the periplatform sediments. In the

case of Sites 1005 and 817, the pore waters always appear to have the same 87Sr/86Sr ratio as

the contemporaneous sediments. This is probably a result of the high rate of carbonate

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dissolution and precipitation relative to the rate of diffusion. In contrast, the behaviour

observed at ODP Site 823 is closer to that observed in the oceanic studies.

Attempts have been made to apply the methods to calculate rates of dissolution and

precipitation outlined by Baker et al. (1982) to periplatform deposits from the Bahamian (Swart

& Guzikowski, 1988) and Maldives carbonate platforms (Swart & Burns, 1990). In both

locations the Sr diffusion method was determined to give low estimates of dissolution and

precipitation. Since these sediments are composed of a significant amount of aragonite it

would be intuitive that these sites would show higher estimated rates of dissolution and

precipitation. It was concluded that the low rates were a result of the high initial Sr content of

the sediments without a corresponding increase in the Sr gradient in the pore waters and were

probably a result of the formation of celestite. While the Sr-diffusion method gave low rates of

recrystallization, the equilibration method was not considered valuable as an indicator of

sediment dissolution and precipitation in situations where there was an initially high

concentration of aragonite or the composition of the sediments was variable through time

(Swart & Burns, 1990).

Oxygen Isotopes

Typically the pore fluids of sediments deposited adjacent to carbonate platforms, such as the

Bahamas, show a steady increase in the 18Ow value of the pore fluids as a result of dissolution

and precipitation of carbonate minerals at increasingly higher temperatures (Swart, 2000).

Figure 22 shows the patterns of 18Ow values in the pore water obtained from five cores drilled

to depths of ca 1000 m adjacent to GBB. The carbonate content of the sediments at all of these

sites is typically in excess of 95% (Eberli et al., 1997). Although at the top of each of the cores

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there is a slight increase in the 18Ow values of the pore fluids related to the increase in the

18Ow values of glacial seawater (Adkins & Schrag, 2003; Schrag, 1996), the steady increase in

18Ow of the fluids with depth reflects dissolution and precipitation at higher temperature and

can be modelled using a similar approach to that outlined for oceanic sediments (Eq. 7). The

results of the 18O values of the bulk sediment, diagenetic component and pore waters are

shown in Fig. 23. Using a rate of recrystallization of 3% per 100 m the models agree with the

measured pore water and sediment data.

While the example of the Bahamas serves to illustrate the differences between oceanic and

platform carbonates, the data only show changes in the upper ca 1000 m below the sea floor.

In fact the Bahamas rests on about 5000 m of mixed carbonates and evaporites which in turn

overlie Jurassic siliciclastics (Goodell & Garman, 1969; Schlager et al., 1988). The extent to

which these sediments are influenced by post depositional/diagenetic process is not known.

However, one process for which evidence exists at the present day is a steady increase in Cl-

with depth in the upper 1000 m (Eberli et al., 1997). Such an increase has also been observed in

other areas associated with carbonate platforms (Davies et al., 1991). Although it has been

suggested that this increase results from dissolution of an underlying NaCl unit and subsequent

upward diffusion, an alternative hypothesis is that the increase is simply a consequence of

continual dissolution and precipitation (Kramer et al., 2000) of carbonate. The Cl- is derived

from the release of this element during dissolution and its subsequent discrimination during

dissolution and precipitation. If the high concentration of Cl- suggests higher amounts of

dissolution and precipitation at greater depths it is also possible that the 18Ow values of pore

fluids could be elevated as high as +8 to +10‰ depending on the amount of dissolution and

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precipitation, burial depth and geothermal gradient. Such increases in the18Ow values occur

without changes in salinity, hence decoupling salinity and 18Ow. Once formed such positive

18Ow fluids can be mobilized through fractures and react with shallow carbonate sediments.

Such fluids are likely to be Ca2+ rich, but Mg2+ poor as a result of dolomite formation and are

not likely, therefore, to be dolomitising, but rather leading to dedolomitising reactions and the

precipitation of LMC.

Oxidation of Organic Material

Oxidation of OM and its subsequent influence upon the pore waters and sediments is generally

considerably more significant in platform-derived sediments than in oceanic carbonates,

principally as a result of the much higher burial rates found at such sites. With increasing depth

the SO42- can become completely exhausted leading to large increases in alkalinity. Other trace

components such as NO3-, NH4+ and PO43- derived from the degradation of OM might also be

visible in the pore waters, although these tend to be utilized by bacteria, absorbed onto

carbonate grains and, in the case of NO3-, used as an electron acceptor during oxidation of OM.

Depending on the extent of the carbonate reactions the alkalinity can remain high and conform

to the 2:1 relationship as expressed in Eq. 4 (Fig. 15), or the alkalinity can be consumed in

carbonate precipitation reactions. During the bacterial consumption of SO42- the 34S values of

the residual pore water increases and CAS incorporated into diagenetic carbonate along this

gradient will have 34S values higher than expected. The alkalinity produced from the oxidation

of OM, while tending to have a 13C value which reflects that of the OM, is usually masked by

the C released from the parent carbonate sediments. Because the absolute amount of SO42- in

the pore water within a given amount of sediment is small and the supply from diffusion

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relatively unimportant unless depositional rates are very slow, SO42- reduction cannot alter the

13C value of the sediment unless the carbonate content is low.

Below the zone of SO42- reduction, OM is oxidized by methanogens using CO2 and producing

CH4. The CH4 can migrate upwards through the sediment column where it can either escape to

the sea floor, be consumed by bacteria, or form CH4 hydrates, usually in association with H2S

and CO2 (Kvenvolden, 1981). The process of hydrate formation alters the salinity of the pore

fluids and produces positive 18Ow values as 16O is preferentially incorporated into the gas

hydrate. This positive 18O value can be recognised in diagenetic carbonates formed in

association with the hydrates (Hesse & Harrison, 1981). In the case of carbonate dominated

sediments, the positive 13C value of the CO2 produced during methanogenesis is not evident in

the bulk 13C value as the amount of carbon produced is small compared to the C in the parent

sediment. For example, during Leg 182 extremely high amounts of biogenic methane were

found, yet there was no noticeable influence on the 13C values of the sediments (Feary et al.,

1998; Swart et al., 2000). This process has been described in ancient carbonate poor sediments

(Irwin et al., 1977).

Implications for Diagenetic Carbonates

With increasing burial depth the 18O values of diagenetic carbonates decreases, reflecting

dissolution and precipitation under the influence of increasing burial temperature. In

sediments which are largely composed of carbonate, the 13C values of the altered sediments

typically will not change as the oxidation potential of pore waters in a diffusive regime is too

low to influence the 13C value of the carbonate (as in the case of the oceanic carbonates).

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While there is a definite trend for increased carbonate alteration with increasing depth, this

trend can be modified by the permeability of the sediments and the amount of siliciclastic

material. For example, in locations off of the Bahamas which show alternations in the amount

of non-carbonate material, the intervals with lower carbonate percentage show better

preservation of the original carbonate constituents than the high-carbonate intervals (Frank &

Bernet, 2000). These intervals also possessed more negative 13C and 18O values, attributes

which could reflect diagenesis or more likely the 13C and 18O values of different source

materials.

The Sr content of the diagenetic carbonates would be expected to be high, reflecting the extent

of sulphate reduction in the sediments. In some instances, calcites with Sr concentrations as

high as 3000 ppm (Dix, 1995) and dolomites with values of 2000 ppm (Swart & Melim, 2000)

have been reported. Such high Sr concentrations in ancient rocks might be misinterpreted as

reflecting precipitation from hypersaline brines. Because the favoured mineralogy of calcium

carbonate precipitated in the oceans has changed through time (Sandberg, 1983), the nature of

periplatform sediments will also have varied. Hence during the Cretaceous, a period of calcite

seas, periplatform sediments would have behaved more like oceanic sediments regarding the

buildup of Sr in the pore fluids. Consequently, it might be expected that the Sr concentration of

diagenetic calcite formed during these times will be lower than during times of aragonite seas.

Trace Elements and Stable Isotope Distribution as Indicators of Fluid Flow

During the movement of fluids through carbonate edifices, the dissolution and precipitation of

carbonates will cause elements with low distribution coefficients for the precipitation of LMC

and dolomite, such as Sr, to increase in concentration in the fluids in the down-flow direction.

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Elements which have distribution coefficients >1 (Mn and Fe) will tend to decrease in

concentration in the same direction. This approach has been proposed to trace the direction of

fluid flow and hence constrain the nature of diagenesis, although the results have been

inconclusive (Vahrenkamp & Swart, 1994). The 18Ow values and diagenetic carbonate formed

as a consequence can also change in the direction of fluid flow. For example, fluids which are

heated then advected upwards will gradually cool meaning that there should be a trend for the

18O values of the altered carbonates to be more positive away from the cooling. Conversely,

fluids being drawn into the margin of a carbonate platform, by a process such as Kohout

convection, will be initially cold and gradually increase in temperature. During this process the

diagenetic carbonate will be isotopically more negative and the fluids more positive, similar to

the case of sediments being buried adjacent to the margin of a carbonate platform.

Examples of Advection in Carbonate Platforms

An example of large scale advection is evident from the pore water and solid geochemistry of

samples from ODP site 812 drilled on the Queensland Plateau, a submerged carbonate

platform, located off of the eastern coast of Australia. Here coring of the upper ca 40 m of

sediments revealed that the pore water profiles were controlled by diffusion showing patterns

typical of sediments deposited adjacent to carbonate platforms (Fig. 24; Davies et al., 1991). At

about 40 mbsf, a hard cemented layer was encountered which was not recovered. Located

below this layer was ca 100 m of unconsolidated sediment containing an abundance of

dolomite sand and silt. The concentrations of elements such as Ca2+, Mg2+, Sr2+ and SO42-within

these sediments were similar to that of seawater (Swart, 1993). The 87Sr/86Sr of the pore fluids

in the lower section also returned values similar to those in modern seawater (Elderfield et al.,

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1993). After the hole was completed a temperature log of the open hole revealed isothermal

conditions from the sea floor to the well bottom, indicating that the formation was under

pressured and seawater was being sucked into the carbonate platform (Davies et al., 1991).

What was encountered at ODP Site 812 was probably part of a large circulation cell involving

modern seawater in the Queensland Plateau perhaps driven by the temperature differences

between the platform and the adjacent seaways (Kohout Convection). In this case the

circulation of seawater probably promoted the dolomitisation of the sediments. The 87Sr/86Sr

of dolomites in these sediments was measured and found to fall between the values typical for

sediments of this age and modern seawater (McKenzie et al., 1993; Fig. 24). Systems such as

this might be responsible for the formation of the precursors of marine fibrous calcites, a

common form of cement found precipitated in the pore space of carbonate accumulations

throughout the Phanerozoic (Kim & Lee, 2003; Richter et al., 2011; Tucker, 2001). Although in

most instances these are composed of LMC [fibrous dolomites have also been reported (Richter

et al., 2014)] such cements have been suggested to be either original (Kendall, 1985; Saller,

1986) or replacements of either aragonite (Kendall & Tucker, 1973) or HMC cements (Lohmann

& Meyers, 1977; Mazzullo et al., 1990; Wilson & Dickson, 1996). The presence of micro-

inclusions of dolomite or empty voids has been used as evidence of a prioir HMC mineralogy,

but these features are not always present making the interpretation controversial. Although

their apparent paucity during the Tertiary is problematic, it is possible that they represent the

replacement of either modern HMC or aragonite cements found in modern reef framework as

described by Aissaoui (1988) and others (Ginsburg & James, 1976; Grammer et al., 1993).

Differences in the original mineralogy of these cements could be induced by changes in the

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Mg/Ca ratio of the ambient environment stemming from other geochemical processes within

the carbonate framework, changes of the Mg/Ca ratio in the oceans and/or changes relating to

the saturation state at the depth of formation. Although these cements have been widely used

as a proxy for interpreting the geochemistry of seawater, it has been suggested they may not

be reliable for this purpose (Wilson & Dickson, 1996) .

Continental Margins Carbonates

The principal reason why continental margins have been considered separately from oceanic

and platform sediments is because of the influence of the hydrological head on the adjacent

mainland. These fluids can be either meteoric in origin, and therefore possess low salinity, or

be remnants of hypersaline fluids. For example, ODP Site 905 drilled along the New Jersey

margin showed a decrease in Cl- with depth suggesting freshwater influence (Gieskes, 1974;

Mountain et al., 1994) probably originating from the US mainland. Similar explanations have

been proposed for ODP sites drilled off the coast of South America (Kastner et al., 1990). In

contrast, other sites show evidence of saline brines (Feary et al., 1998; Kastner et al., 1990;

Sayles & Manheim, 1975); for example, sites drilled during ODP leg 182 off of the southern

coast of Australia in the Great Australian Bight (GAB). These sites exhibited large increases in

salinity with increasing depth suggesting the presence of hypersaline lagoons on the adjacent

mainland during previous sea-level low stands (Feary et al., 1998; Fig. 25). In addition to the

role of non-marine fluids, sediments along continental margins tend to be principally siliciclastic

and are therefore relatively carbonate poor and rich in OM. For these reasons the

concentrations of Sr2+ in the interstitial pore fluids tends to be low in spite of abundant SO42-

reduction. The 87Sr/86Sr profiles at such sites are complicated, reflecting the input of Sr both

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from the dissolution and precipitation of any carbonate present, as well as alteration of any

felsic minerals which typically have very elevated 87Sr/86Sr ratios (Kastner et al., 1990).

Concentrations of Ca2+ and Mg2+ tend to be low in settings unaffected by saline brines as

authigenic carbonates are precipitated consuming available Ca2+, Mg2+ and CO32-. In the

presence of saline brines the concentrations of all interstitial components can rise depending

upon the rates of reactions. For example, at Sites 681 and 680 off the Peruvian margin,

concentrations of Ca2+ and Mg2+ increased dramatically with depth and behaved more or less

conservatively (Kastner et al., 1990). In contrast, at sites off of the GAB (1127, 1129 and 1131),

concentrations of Ca2+ and Mg2+ stayed low despite a threefold increase in salinity (Feary et al.,

1998). Here both Ca2+ and Mg2+ are being consumed in carbonate precipitation reactions. An

example from Site 1127 is shown in Fig. 21B. Typically 18Ow values of pore fluids found in

continental sediments are negative reflecting mineral–rock interactions with silicates as well as

possible meteoric sources.

Oxidation of Organic Material

Many continental margin settings tend to have high rates of upwelling and hence high organic

productivity (Emeis & Morse, 1990; Emeis et al., 1991). In such locations sediments tend to be

siliceous, carbonate poor and rich in OM. The high organic content arises as OM is produced,

not only on adjacent shelf deposits and in the water column but also from rivers draining the

mainland. Typically, therefore, both the rates of preservation and oxidation of OM are high

and the pore fluids significantly altered with respect to the concentration of SO42-, alkalinity and

other constituents associated with the degradation of OM (Kastner et al., 1990). Organic

material frequently survives below the zone of SO42- reduction and methanogenesis produces

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copious amounts of CH4. As a consequence, gas hydrates tend to be more abundant along

continental margins compared to other locations. Continental margins also tend to have lower

carbonate contributions as a result of dilution by siliciclastic material. Consequently diagenetic

carbonates formed in these environments could conceivably be influenced by the negative 13C

value of the CO2 produced during SO42- reduction and methanogenesis. The 13C values of

carbonates found in such locations are therefore not related to the 13C value of the original

carbonates or the 13C value of the DIC in the ocean during deposition and the use of 13C

values from such sections to interpret changes in the global 13C cycle should be avoided. In

instances where pore fluids are influenced by adjacent saline brines an additional supply of

SO42- is provided, allowing greater amounts of OM to be oxidized and higher alkalinity pore

waters to be produced. In the example from Leg 182 off the GAB, the three-fold increase in

salinity produced a similar initial increase in SO42-. This SO42- was consumed during sulphate

reduction producing alkalinity values over 100 mM. Such high amounts of sulphate reduction

produced high concentrations of H2S which, together with CH4 formed by methanogenesis,

resulted in the possible production of CH4–H2S–CO2 hydrates at this location (Swart et al.,

2000). However, because the sediments were composed of >90% carbonate minerals, even this

extreme degree of SO42- reduction was insufficient to alter the bulk 13C value of the sediment.

Implications for Diagenetic Carbonates

The geochemistry of diagenetic carbonates in settings influenced by continental fluids can be

quite variable because in such situations both meteoric and hypersaline fluids have been shown

to be present. In instances where there are low initial amounts of carbonate sediments, the

13C value of the pore water can become dominated by the products of the oxidation of OM

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and therefore any diagenetic cement will be isotopically negative. In carbonate dominated

sediments such as those occurring off the GAB, even extensive OM diagenesis does not

significantly influence the 13C value of the sediments.

Deeper Burial Diagenesis

While sediments deposited above oceanic crust experience limited burial, carbonates formed

along continental margins and on carbonate platforms can be eventually buried to great

depths. Therefore at depths of 2 to 3 km, where formation temperatures can be 60 to 100oC,

dissolution and precipitation reactions are enhanced. Where fluids have encountered

evaporites they may be Na rich and have very high salt contents (Land & Prezbindowski, 1981).

Continued movement of such fluids results in the formation of diagenetic evaporate minerals

displaced from their original depositional sites (Machel, 1993). Carbonate recrystallization at

high temperatures can elevate the fluid 18O and in contrast to the example from the Bahamas

where 18O values reach ca 2 to 3‰ (Fig. 19), formation water 18Ow values of over +20‰ have

been reported for the Cretaceous (Land & Prezbindowski, 1981). Fluids originating at these

depths can be mobilized into shallower environments along faults, allowing hydrothermal

processes to affect rocks in much shallower burial environments. Although typical 18O values

of such hydrothermally altered rocks are very negative (ca -10 to -20‰; Katz et al., 2006),

suggesting high temperatures, as the 18O value is a product of both temperature and the

18Ow, the interpretation is often complex and even carbonates with more positive 18O values

might form at high temperatures.

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Hydrothermal Dolomitisation

Hydrothermal dolomites are defined as those formed at temperatures higher than the ambient

formation temperatures (Davies & Smith Jr., 2006) and are therefore associated with fault

systems controlled by structural deformation. Such faults allow higher temperature fluids to

migrate from deeper formations and interact with the host rocks precipitating and/or dissolving

carbonates along the way. For example, such a mechanism has been proposed for the

formation of the dolomites in the Arab-D Formation (Swart et al., 2005) and has recently been

supported by the use of clumped isotopes which documented that the dolomites formed at

temperatures in excess of 60 to 70oC and from fluids with positive 18Ow values (Swart et al.,

2013). Such deep basinal fluids, however, are probably not primarily dolomitising being

depleted in Ca2+ and Mg2+ and probably corrosive to carbonate minerals; they can therefore

dissolve the calcite from partially dolomitised bodies creating a highly permeable and porous

reservoir (Montanez, 1994; Swart et al., 2005).

Stylolites

Stylolites are roughly planar features, present at a variety of scales from microscopic to metre

length or longer, which represent surfaces where rock material has been lost by pressure

solution. During compaction there is typically dissolution of the host carbonate, a loss of

porosity (Rittenhouse, 1971) and insoluble material (clay, OM, pyrite, etc.) is left behind leaving

a distinct planar or ragged surface. Stylolites are of particular interest in the study of reservoir

characterisation because they act as permeability barriers (Heap et al., 2014); they have been

extensively studied (Rye & Bradbury, 1988; Saller & Dickson, 2011; Swennen et al., 2012) .

Some work shows isotopically more negative 13C and 18O values associated with the stylolites

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(Rye & Bradbury, 1988), while no changes were observed in others (Hilgers et al., 2006). The

changes in the 18O values can be interpreted as dissolution of carbonate and reprecipitation at

higher temperatures, while the more negative 13C values indicate the incorporation of

isotopically negative carbon perhaps released during the thermal degradation of OM.

Thermo-Chemical Sulphate Reduction

At high temperatures (120 to 140oC), evaporite minerals (gypsum and/or anhydrite) and/or

sulphate within the formation can react with any petroleum present to produce significant

quantities of H2S and H2O according to a simple reaction as shown in Eq. 8 (Vandeginste et al.,

2009; Worden et al., 2004; Worden & Smalley, 1996). This process is known as thermo-

chemical sulphate reduction (TSR) with a typical example shown in Fig. 26 where the reaction of

anhydrite with hydrocarbon forms calcite and pyrite:

(8)

This process results in the formation of isotopically light carbonate, because the 13C value is

inherited from the CH4 and more negative 18Ow values (Worden & Smalley, 1996). The salinity

of fluids associated with TSR is substantially lower than the normal formation fluids. Typically

the diagenetic carbonate forms rinds around existing evaporite minerals with relatively

negative 13C and 18O values (Vandeginste et al., 2009). In some instances pyrite and

elemental sulphur are formed in the process. Based on the geochemical signatures and mineral

associations, some authors have invoked TSR related reactions during the formation of saddle

dolomite (Machel, 1987b), a type of dolomite known to be formed at high temperatures (Radke

& Mathis, 1980).

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Chemostratigraphy and Diagenesis

Chemostratigraphy in carbonate rocks involves recognising distinctive patterns in the vertical

profile of a geochemical parameter and then assigning ages to specific inflections in the record

or correlating the inflections between locations with similar profiles. Often the inflection points

have not actually been dated and it is only assumed that similar patterns at geographically

disparate locations represent the same phenomenon. Frequently used geochemical

stratigraphic tools include 13C, 18O, 34S and the 87Sr/86Sr ratio (Veizer et al., 1999).

Strontium Isotopes

The 87Sr/86Sr of marine carbonates has changed gradually over time reflecting a balance

between the erosion of continental rocks, which typically have high 87Sr/86Sr ratios, and the

recycling of ocean water through mid-oceanic ridges where the 87Sr/86Sr ratio is low (Burke et

al., 1982). In oceanic settings as described earlier, diffusion of Sr along a concentration gradient

in the pore fluids can change the 87Sr/86Sr of the sediments if there is a significant amount of

alteration. If the rates of dissolution and precipitation reactions are high, as in the case of

platform and periplatform carbonates, then the 87Sr/86Sr of the pore fluids tends to be

indistinguishable from that of the original marine carbonates. Diagenetic carbonates formed

along such a profile have 87Sr/86Sr ratios similar to the original sediments and therefore, in spite

of diagenesis (Swart et al., 2001), this ratio can be used to obtain some idea of the depositional

age by comparing the ratio in the rock with established 87Sr/86Sr curves for the oceans (DePaolo,

1985; Koepnick et al., 1985; McArthur, 1994). In other instances, the 87Sr/86Sr ratio can be used

constrain the timing of diagenesis provided the Sr source to the system is known (Saller, 1984;

Saller & Koepnick, 1990).

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Oxygen Isotopes

Variations in the18O values of sediments have proved to be important for dating during the

Pleistocene (Emiliani, 1955; Lisiecki & Raymo, 2005). The 18Ow values of the oceans increased

during glacial periods, reflecting the buildup of continental ice, and then subsequently

decreased as ice melted. By analyzing organisms such as foraminifera which continually track

these changes, the age of sediments with maximum or minimum 18O values can be dated by

reference to known 18O curves. By comparing the variations of 18O values in benthonic and

planktonic foraminifera, the relative influence of temperature and ice volume has been

ascertained. During the Holocene and Pleistocene, the large temperature difference between

surface and deeper waters is reflected in the 18O values of benthonic and planktonic

foraminifera. This difference decreases as one progresses back in time and has been

interpreted as indicating that deep waters were warmer in the past (Savin, 1977). An

alternative interpretation is that the smaller numbers of benthonic foraminifera have been

affected by the dissolution and precipitation reactions of the more abundant planktonic

foraminifera. With time, dissolution and precipitation in a closed system will homogenise the

18O values of benthonic and planktonic foraminifera. Any differences between the 18O values

of the benthonic and planktonic foraminifera, therefore, might in fact be a diagenetic artifact

rather than be indicative of changing bottom water temperatures (Killingley, 1983).

In sediments older than the Miocene, diagenesis is too pervasive to allow the 18O value of

carbonates to be useful for detailed chronostratigraphy. While there is a tendency for the 18O

values of carbonates, marine cherts and phosphorites to become increasingly negative over the

past 600 Myr (Degens & Epstein, 1964; Keith & Weber, 1964; Knauth & Lowe, 1978; Veizer &

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Hoefs, 1976), the causative mechanism is debatable. Speculation has centred on increased

temperature of the early Earth, increased diagenetic alteration, or changing 18Ow values of the

oceans. The various theories have been presented in numerous papers (Knauth & Epstein,

1976; Knauth & Lowe, 1978; Perry, 1967; Perry & Tan, 1972) and new life has been injected into

the debate with the emergence of the clumped isotope technique which theoretically should be

able to distinguish between the various options, providing that solid state resetting has not

taken place.

Carbon Isotopes

The 13C values of carbonate sediments in conjunction with the 13C values of OM have been

used extensively to document changes in the global carbon cycle (Hayes et al., 1999). Increases

in 13C values have been mainly interpreted as reflecting an increase in the burial of OM, while

decreases indicate increased oxidation of OM. Correlations between the 13C value of inorganic

and OM is usually taken to indicate a 13C signal unaltered by diagenesis, and therefore one

which reflects perturbations in the global carbon cycle (Bachan et al., 2012). Many detailed

13C records have been measured, principally in bulk carbonate rocks, and these have been

correlated globally for large portions of the Phanerozoic and Proterozoic rock record (Kennedy,

1996; Koch et al., 2014; McKirdy et al., 2001; Saltzman et al., 2004). Although the best samples

for determining global changes in the carbon cycle are pelagic carbonates such as foraminifera,

or even bulk oceanic sediments, these are not available for time periods older than the

Cretaceous–Jurassic when records from carbonate platforms or epeiric seas have to be used.

There are problems with such locations and widespread disagreement as to whether these

records reflect differences in sediment types, local or global variations in the carbon cycle or

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diagenetic alteration (Derry, 2010; Holmden et al., 1998; Knoll et al., 1986; Swart & Eberli,

2005). Despite an almost euphoric tendency to accept that global changes in the 13C values of

carbonates are original in nature, there are several viable alternative explanations for these

global trends including variations in the origin of the sedimentary materials and diagenesis.

Origin of Sediments

The origin of carbonate sediments and their influence on 13C stratigraphy and interpretations

of the global carbon cycle can be seen in a study of sediments deposited along the slope of the

GBB (Swart & Eberli, 2005). Such a setting might be analogous to an ancient carbonate margin,

although the types of organisms and their mineralogy have undoubtedly changed through time.

At this location a series of five cores were taken during ODP Leg 166 (Eberli et al., 1997;

Shipboard Scientific Party, 1997; Fig. 12). In previous studies it was shown that the composition

of periplatform sediments respond to changes in sea level (Droxler et al., 1983; Reuning et al.,

2006; Roth & Reijmer, 2005). As the platform surface material is isotopically enriched in 13C,

the 13C signal in periplatform sediments also varies as a function of sea-level change (Swart &

Eberli, 2005). During highstands, abundant shallow-water material is produced on the platform

top and deposited along the margins. During lowstands, the principle source of carbonate is

pelagic organisms. Because these two sources have quite different 13C values, the result is a

correlation between sea level and 13C values, as well as a correlation between the 13C value

and position relative to the platform margin. Using bio-stratigraphy and seismic stratigraphy,

the 13C changes of the bulk carbonate were correlated between the five cores in a proximal to

distal transect. In an ancient setting such changes in 13C values might be interpreted as

reflecting changes in the global carbon cycle, but in the case of the Bahamas, they do not

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correlate to the well-established decrease in the 13C value of the oceans (Shackleton, 1985).

There is an absence of correlation because changes in the Bahamas are driven by global sea

level, not the global carbon cycle, and in fact similar changes in 13C values have been observed

along carbonate platform margins in the Indian and Pacific Oceans (Swart, 2008). Another

feature of the sediments at these locations is that their 13C values become more positive

towards the present day. This trend was interpreted to reflect a gradual progradation of the

platform margins towards the present position of the cores, reflecting increased input of

isotopically positive material produced on the platform top.

Carbonate Diagenesis: Inorganic carbon

The potential influence of diagenesis upon the 13C value of shallow water carbonates is well-

established. Consider the example of the Bahamas transect cores in which the 13C values of

individual sequences from ODP Sites 1004, 1005, 1006 and 1007 fall approximately on a 1:1 line

(Swart & Eberli, 2005) when plotted against data from Site 1003. The data from cores Clino and

Unda, which were cored on the surface of GBB, an area subject to repeated sea-level changes,

clearly do not plot on this line (Fig. 27). Curiously the types of trends in 13C and 18O values

associated with Pleistocene sea-level changes (Allan & Matthews, 1977; Melim et al., 2004;

Melim et al., 2001; Melim et al., 2002) are rare throughout large periods of the geological

record. In fact, where similar magnitude changes do exist such as during the Neoproterozoic or

the Permian–Triassic extinction (Krull et al., 2004), they have been mainly (but not exclusively)

interpreted as reflecting changes in the global carbon cycle (Hoffman et al., 1998; Macdonald et

al., 2010). As discussed earlier, it is difficult for samples which have experienced only marine

diagenesis to become significantly altered from their original 13C composition. Therefore

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carbonates deposited below the range of depths affected by glacial-eustatic variations in sea

level are preferable for chronostratigraphy because they avoid most diagenetic effects. High

temperature and thermal decarboxylation of OM have, however, also been suggested as an

alternative mechanism for altering the 13C value of carbonates (Derry, 2010). Despite this

preference for deep-water deposits in chemostratigraphy, there have been numerous examples

of shallow water carbonates apparently unaltered by freshwater in that they appear to have

relatively positive 13C values which can be correlated over wide geographic areas (Saltzman,

2002; Saltzman et al., 2004; Vahrenkamp, 1996). It seems likely that there were only minor

changes in sea level during deposition of these samples, and therefore exposure to freshwater

was limited or arid climates prevailed. Such small changes in sea level still produce variation in

the 13C record (Immenhauser et al., 2003; Immenhauser et al., 2002; Rameil et al., 2012), but

the effects are more subtle than those of the Pleistocene and are often confused with changes

produced by secular variations in the global carbon cycle.

Alteration of sediments near or at marine hardgrounds and/or in sediments with reduced

deposition can allow the 13C values of the carbonate to become altered, usually towards more

negative values as 12CO2 is released by respiration and sufficient oxidants are supplied through

diffusion. Such signals could conceivably be associated with well-known events in Earth history,

such as the Palaeocene–Eocene boundary, and might enhance oceanic derived 13C signals in

the sediments overestimating the amounts of OM requiring oxidization, or the processes

needed to produce global changes in the oceanic 13C value. Furthermore, marine alteration of

the 13C signal can take place in low-carbonate sediments because the pore water will no

longer be buffered to the same extent.

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Variations in the 13C values of carbonate have recently been explained by a different

mechanism (Schrag et al., 2013). This idea suggests that significant authigenic carbonate with

isotopically negative 13C values derived from anoxic oxidation of organic matter formed at or

near the sea floor. Hence changes in the 13C values of carbonates would not necessarily be

coupled to the burial of organic carbon during periods of ocean anoxia.

Carbonate Diagenesis: Organic Carbon

Mixing as a source of variation

Because the 13C value of OM is often used to support evidence for changes in the organic

carbon cycle, Oehlert et al. (2012) investigated how the 13C value of the OM was related to the

13C value of the inorganic fraction of the ODP Leg 166 samples as measured by Swart & Eberli

(2005). Oehlert et al. (2012) showed a variable degree of correlation between the 13C values

of the inorganic and organic fraction depending upon the position of the core relative to the

platform margin. Closer to the margin of GBB (ODP Site 1005), where the majority of the OM is

derived from shallow water carbonate organisms such as sea grasses, there was no correlation

between the 13C values of the organic and inorganic fractions. Further away from the margin

the degree of correlation improved as OM with more positive 13C values originating from the

platform top mixed with material with relatively more negative 13C values derived from pelagic

sources (Site 1006). This change in the correlation is once again a feature of mixing between

platform derived materials with relatively positive 13C values in organic and inorganic

components, with isotopically lighter pelagic material. Closer to the platform the correlations

between the inorganic and organic fractions breaks down because the sediments contain a

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larger proportion of platform derived material. As the surface sediments themselves have

quite a wide range of 13C values in the OM, but a relatively narrow range of 13C values in the

inorganic fraction, there is no correlation between the 13C value of inorganic material and that

of OM on the platform surface and consequently this is reflected at the proximal sites.

Diagenesis as a source of variation

It has been stated that positive correlations between the 13C value of organic and inorganic

carbon in the geological record must reflect changes in the original carbon cycle. Specifically

Knoll et al. (1986) says that: “…no secondary processes are known (or for that matter

conceivable) which always shift the isotopic composition of carbonate and organic carbon in

the same direction at the same rate”. However, not only do correlations between the 13C

values of organic and inorganic carbon arise as a result of the mixing of OM in the depositional

environment, as shown by Oehlert et al. (2012), but different diagenetic processes operating at

the same time can result in processes which produce strong correlations between the 13C of

organic and inorganic carbon. For example, in altered Bahamian carbonates, a strong positive

correlation exists between the 13C value of organic and inorganic carbon (Oehlert & Swart,

2014; Fig. 28). This strong correlation clearly does not mean that the changes are related to the

global carbon cycle. In fact what takes place during sub-aerial exposure is that the carbonates

are altered by meteoric fluids in which the 13C value of the DIC becomes more negative as a

result of the addition of CO2 derived from the oxidation of OM. At the same time, OM from

terrestrial vegetation and algae is added to the semi-consolidated sediment which is in the

process of becoming lithified. In the case of exposed carbonates, such as the Bahamas, this OM

is mainly derived from C3 plants which are approximately 10 to 15‰ more negative than OM

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found on the surface of the shallow-carbonate platform. Hence, diagenetic processes

associated with sub-aerial exposure result in both the OM and the carbonate acquiring an

isotopically more negative 13C signal than the original sediment. When a section or core is

taken through a sedimentary sequence which has been altered in this manner, the result would

be an apparent strong positive correlation between the 13C values of organic and inorganic

carbon, a trend which has been interpreted in numerous publications as being original (Knoll et

al., 1986; Rose et al., 2012). While it would not be appropriate to say that all sections which

show a positive correlation between the 13C value of organic and inorganic carbonate are a

result of diagenesis, neither would it be correct to state that there are no secondary processes

which could conceivably alter the 13C value of organic and inorganic carbon in the same

direction.

Sulphur Isotopes

The 34S value of the oceans as measured in evaporite minerals (Claypool et al., 1980), CAS (Gill

et al., 2011) and barite (Griffith & Paytan, 2012) has revealed a secular change in the 34S values

of the oceans. These changes arise as a result of a shift in sulphur between the various

reservoirs in the oceans, evaporites and inorganic minerals such as sulphide bearing minerals.

Over longer time scales the 13C values of unaltered carbonates and the 34S value of evaporites

tend to be inversely correlated (Veizer et al., 1980). Such a correlation may arise because

during periods of increased photosynthesis, leading to increased burial of OM, and hence

during which the 13C values of the carbonate record are more positive oxygen levels in the

atmosphere are higher. This leads to higher rates of weathering of sulphide bearing minerals

which normally tend to have more negative 34S values. However, positive correlations

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between the 34S of the CAS and 13C values of the carbonate and OM can also occur associated

with oceanic anoxic events (OAEs), such as those of the Cretaceous (Jenkyns, 2010). Under

such circumstances, the positive correlation between 13C and 34S values results from

widespread anoxia in the deeper water column promoted by high rates of organic production,

causing isotopic enrichment in 13C of carbonate formed in the surface oceans. The OM

descending into the deeper ocean caused partial sulphate reduction leading to a residual pool

of sulphate which is isotopically enriched in 34S. This pool is partially upwelled to the surface

oceans returning a positive source of 34S to be subsequently incorporated into the CAS.

One concern about the 34S values of CAS is its integrity during diagenesis. While the 34S

values of carbonates altered during freshwater diagenesis are retained (Gill et al., 2008), during

marine burial the concentration of SO42- in pore water is frequently completely exhausted

during oxidation of OM. Because there is significant isotopic fractionation during bacterial

sulphate reduction (BSR), pore waters become greatly enriched in 34S. Sulphate with these

elevated 34S values can become incorporated into the CAS during carbonate transformation

reactions. As a consequence, the veracity of the 34S signal in CAS from sections experiencing

carbonate dissolution and precipitation, accompanied by sulphate reduction, must be

questioned. The situation can be simply modelled by examining what would happen to the 34S

values of the bulk carbonate as it is altered in a sedimentary sequence which experiences

complete loss of the interstitial sulphate in the pore water as a consequence of BSR. Along this

same pathway the original biogenic sediments are dissolved and inorganic LMC and dolomite

precipitate. Such dissolution and precipitation is common in the majority of DSDP and ODP

sites (Baker, 1985; Gieskes, 1976; Kastner et al., 1990; Kramer et al., 2000). During BSR, 32S is

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preferentially incorporated in the H2S leaving the residual sulphate pool enriched in 34S. If it is

assumed that 100% of the carbonate is altered to LMC by the time the sulphate is consumed

with a value of 1.04, then each increment of carbonate which is precipitated will have a

progressively more positive 34S value reflecting the modified pore fluid. Using modern values

for oceanic SO42- concentrations (28 mM) and an initial concentration of S in the carbonate of

4000 ppm, then the 34S value of the LMC at the bottom of the sulphate reduction zone would

increase to +29‰, 9‰ more positive than the present oceanic 34S value (Fig. 29). If the SO42-

concentration of the initial carbonate was lower or the  value >1.04, then the diagenetic effect

would be greater.

Although the above model indicates that there might be significant potential for alteration of

the 34S value of the CAS signal as a consequence of marine diagenesis, the scenario outlined

above would be a worse case situation, as the model predicts continual carbonate dissolution

and precipitation through the zone of sulphate reduction. If the carbonate was completely

altered to LMC early before BSR took place, then the original 34S value of the CAS would be

preserved. Conversely, if the carbonate survived unaltered below the BSR zone, then the 34S

value of the CAS would also be unaffected. Intermediate situations would also result in a

reduced diagenetic effect upon the 34S values of the CAS.

Although the present concentration of sulphate in the oceans is probably as high as it has ever

been (Demicco et al., 2005), lower initial concentrations would not result in a reduced

diagenetic effect because presumably the initial amount of sulphate in the original carbonate

would also have been reduced. Consider the Carboniferous where the SO42 concentration of

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the oceans was approximately 10 mM or 31% of modern values. As a consequence of the lower

initial SO42 the CAS content of the carbonates precipitated in these oceans would also be 30% of

modern values and therefore the model diagenetic effect would be approximately the same.

FINAL CAUTIONARY THOUGHTS

The study of diagenesis in carbonate rocks using geochemistry has reached an exciting point.

Not only have the classic tools of C and O isotopes and trace elements continued to improve in

terms of their accuracy, sensitivity and microsampling ability, but there are a number of

additional geochemical proxies and methods of applying these proxies potentially available.

Some of these proxies need to be refined and there are still other proxies waiting to be

discovered. However, one concern is the rush to apply all geochemical techniques to the

interpretation of ancient rocks without adequately understanding the factors controlling the

proxy. The extensive literature documenting diagenetic changes in modern and recent

sediments and rocks would also repay critical perusal. The geochemical signatures of

carbonates are undoubtedly related to the temperature and geochemical conditions of the

environment in which they formed or in which they were altered. This has been confirmed

through numerous studies where carbonates or calcareous organisms were precipitated/grown

in the laboratory under controlled conditions. The next logical step of such experiments is to

measure geochemical variation in naturally occurring samples. However, application of such

studies, even to very recently formed carbonates, shows that it is difficult to obtain a clean

environmental signal. This is because, in spite of the large amount of research already

performed, there are still many aspects of carbonate formation (skeletal and non-skeletal)

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which are not understood. If it is problematic to interpret proxies in Modern environments, can

there be much faith in interpreting the same proxy in 600 Myr old samples? The literature is

replete with cases where geochemical interpretations have been used to support changes in

the original depositional environment rather than diagenetic changes. For example, it was

recognised very early in the application of stable C and O isotopes to carbonates that diagenesis

plays a critical role in governing the ultimate stable isotope composition of the rock. However,

C and O isotopes are frequently applied to rocks which are hundreds to thousands of millions of

years old without recognizing the basic principles affecting the alteration of sediments to

lithified rocks. The same criticism can be levelled at the numerous other geochemical proxies

applied to ancient rocks. Unfortunately it is more ‘news worthy’ to publish studies which

interpret large changes in the geochemical record as extreme events in Earth’s early history,

rather than as a result of more mundane geochemical process such as diagenesis. If the

present or the Neogene is any kind of key to the past, geochemical studies of ancient rocks

should be rooted in rigorous investigations of modern and near modern environments

combining petrographic and geochemical approaches. So that we do not: “Throw the Baby out

with the Bath Water” (Marshall, 1992), such proxies should be cautiously applied to older

materials. Only then can the potential of geochemical indices be made available to correctly

study both the palaeoenvironment and the diagenesis of carbonates in older time periods.

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ACKNOWLEDGMENTS

I would like to thank my graduate students and post-doctoral associates who have worked on

the meaning of trace elements and stable isotopes in carbonates. These include Monica

Arienzo, Steve Burns, Gong Chung, Mike Guzikowski, Genny Healy, Yula Hernawati, Phil Kramer,

Mike McClain, Sevag Mehterian, Leslie Melim, Sean Murray, Amanda Oehlert, Brad Rosenheim,

Victor Rossinsky, Philip Staudigel, Volker Vahrenkamp and Amanda Waite. Discussions with

David Budd, Tracy Frank, Robert Ginsburg, Kacey Lohmann, Martin Kennedy and Gene Shinn

helped with the preparation of this paper. Victor Rossinsky and Leslie Melim are thanked for

the images in Figs 10 and 19. The paper benefited from reviews by Art Saller, Jim Hendry,

Maurice Tucker and Isabel Montanez. Ali Pourmand commented on the section of REEs. Greta

Mackenzie is thanked for drafting Fig. 9 and final reading of the manuscript.

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Figure Captions

1- The distribution of 18O values in precipitation as calculated using the IAEA database

(Yurtsever, 1975; Yurtsever & Gat, 1981) and using known relationships between 18O

values and temperature elevation, and distance from coastlines (Bowen & Wilkinson,

2002).

2- (A) Changes in the abundance of carbon species comprising the DIC. (B) Changes in the

13C values of the different species contributing to the DIC assuming that the system

maintains a constant 13C value of the DIC of -5‰. (C) Changes of the 18O value of the

sum of the DIC species as a function of pH as calculated using the approach of Zeebe

(Zeebe, 2007; Zeebe & Wolf-Gladrow, 2001). Fractionation factors are taken from Beck

et al. (2005).

3- Solutions for a range of 18O values for fluids used in the precipitation of LMC (dashed

lines; Kim & O'Neil, 1997) and dolomite (solid lines; Sheppard & Schwarcz, 1970). A LMC

with a measured 18O value of approximately -5‰ could form from fluids with a 18O

value of 0‰ at ca 40oC or +4‰ at ca 65oC. A dolomite with the same 18O value would

need a fluid with 18O values between -4‰ and 0‰ to form at the same temperatures.

4- The Meteoric Water Line (MWL) and the pathway taken during the evaporation of fluids

using the model of Gonfiantini (1986) at different relative humilities As a result of

changes in the activities of H and O with increasing ionic strength, the 18O and 2H

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 values of the evaporating solution actually decrease at high evaporation. Such

behaviour does not happen during the evaporation of freshwater.

5- Compilation of 18O and 13C values in aragonite cements from Bermuda and Belize,

(Gonzalez & Lohmann, 1985)❶, the Bahamas (Grammer et al., 1993)❷, and HMC

cements from the Pacific (Aissaoui, 1988)❸, HMC cement from Bermuda, and Belize

(Gonzalez & Lohmann, 1985)❹, Bahamian sediments (Swart et al., 2009)❺, Heron

Island sediment (Weber & Woodhead, 1969)❻, Enewetak sediment (Weber & Schmalz,

1968)❼ and zooxanthellate corals (Swart et al., 1996)❽. The boxes represent

approximate estimates of equilibrium of aragonite, LMC, and HMC deposited at the

surface of Great Bahama Bank.

6- Compilation of the 18O and 13C values of various types of sediments and corals (from

Fig. 5), deep-sea corals (Emiliani et al., 1978; Land et al., 1977), brachiopods (Carpenter

& Lohmann, 1995), calcareous algae (Lowenstam & Epstein, 1957), brachiopods and

echinoderms (Milliman, 1974). The boxes represent approximate estimates of

equilibrium of aragonite, LMC and HMC deposited at the surface of Great Bahama Bank.

7- (A) Photomicrograph of boring algae and fungi using fluorescence microscopy in the

skeleton of a recently deceased coral (scale bar = 500 m). (B) Etch transverse section

through the septa of a coral showing boring organisms penetrating the centres of the

trabecular axis (scale bar = 100 m). (C) SEM of an etched thin section of a recently

deceased coral showing the abundance of boring endolithic organisms (scale bar = 100

m).

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 8- The 13C and 18O values of various dolomites from the Jurassic Arab-D (Swart et al.,

2005), Pliocene San Salvador (Dawans & Swart, 1988) and Clino (Swart & Melim, 2000)

and the Modern Abu Dhabi (McKenzie, 1981).

9- Schematic figure showing the distribution of the different diagenetic zones associated

with early freshwater diagenesis. The pattern of 13C and 18O values is generalized

from the work of Allan & Matthews (1982) as well as this paper.

10- (A) Host Pleistocene aeolinite from Caicos (Bahamas) capped by a thin calcrete

(Rossinsky, 1986). (B) Close up of calcrete in (A) (Image from Rossinsky, 1986). (C)

Peloids from Caicos in the Bahamas cemented by vadose cements and penetrated by

roots from overlying vegetation (Image from Rossinsky, 1986) (scale bar = 500 m). (D)

SEM of vadose cement from Ocean Bight (Bahamas) (Image from McClain et al., 1992).

(E) Freshwater cements and syntaxial overgrowths on an echinoderm fragment from

Clino core at a depth of 131 mbmp. Sample shows dissolution of precursor grains and

preservations of micritic envelopes (scale bar = 500 m). (F) Sample showing the

influence of freshwater alteration of a mudstone in the phreatic zone of the Clino core

(95 mbmp). Although the sample does not necessarily show classic petrographic

indicators of freshwater diagenesis, the 13C and 18O values indicate that this sample

has been affected by freshwater diagenesis (scale bar = 500 m).

11- Data from diagenetically altered carbonate rocks from the Bahamas show the inverted

‘J’ pattern (Lohmann, 1987). Data are from Rossinsky & Swart (1993) and Rossinsky et

al. (1992).

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 12- Map showing the location of the ODP Leg 166 cores (1003-1007), Clino and Unda cores

(Ginsburg, 2001) and Florida Bay (location of Jimmy and Crane Keys).

13- The 18O and 13C values of carbonates (Melim et al., 2001) from a core drilled in the

Bahamas together with the interpretation from the work of Allan & Matthews (1982). In

this core the upper portion was unconsolidated and lost during the drilling. The 13C

values of surface sediments average +4.5‰, but can be as positive as +6‰. Below this

the following zones can be recognised: (1) the vadose zone: The 13C values are highly

variable, while 18O values are negative but more homogenous. Exposure surfaces have

very negative 13C values; (2) the freshwater phreatic zone is characterized by less

variable 13C but similar 18O values to those in the vadose zone; (3) the mixing-zone is a

transition zone between freshwater altered sediments and non-altered marine

sediments. More dissolution than precipitation occurs in this zone and 13C and 18O

values co-vary; (4) the marine phreatic zone is characterized by limited changes in both

13C and 18O values. Variations which do exist such as in the example here relate to the

input of different types of sediment. In this case bank top non-skeletal sediments form

the majority of the sediment in this interval. The two inserts (367 m and 536 m) show

two non-depositional surfaces where sea floor diagenesis can alter the signal of the

sediments as a result of oxidation of OM.

14- Excesses and deficits of non-conservative elements relative to Cl- at ODP Site 1005 (Fig.

9). Deficits are caused by utilization of the element in reactions such as the oxidation of

organic material which utilizes SO42- and produces alkalinity (Fig. 13).

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 15- Relationship between SO42- and alkalinity from data shown in Fig. 11. The slope of the

relationship is approximately two as predicted from Eq. 4. Deviations from this

relationship are caused by precipitation of carbonate minerals.

16- Relationship between Sr2+/Ca2+ and Ca2+/Cl- from ODP Site 1005 (solid circles; Kramer et

al., 2000) and in a core taken from Florida Bay (red diamonds; Swart et al., 1987a; see

Fig. 9 for location). At Site 1005 the change in ratios down core starting in the circle is

represented by the dashed line. The circle represents the approximate starting

composition of seawater. In the upper portion of the sediment column, precipitation of

LMC and dolomite drives the Ca2+/Cl- ratio down and the Sr2+/Ca2+ ratio up. Calcium ions

diffuse downward from the seawater–sediment interface into the Ca2+ minimum zone

as well upwards from the underlying pore waters where there is dissolution of CaCO3.

17- Relationship between Sr2+/Ca2+ and Mg2+/Cl- from ODP Site 1005 (solid circles; Kramer et

al., 2000) and in a core taken from Florida Bay (red diamonds; Swart et al., 1987a; see

Fig. 4 for location). The circle represents the approximate position of seawater. In Crane

Key dissolution of aragonite occurs as there is an increase in Mg2+/Cl- and Ca2+/Cl- ratio

without a change in the Sr2+/Ca2+ (as the distribution coefficient for aragonite is close to

one). These changes suggest a combination of dissolution of aragonite and HMC

combined with the precipitation of LMC. At Site 1005 the change in ratios is represented

by the dashed line. In the upper portion of the sediment column, precipitation of LMC

and dolomite drives the Mg2+/Cl- ratio down and the Sr2+/Ca2+ ratio up.

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18- Changes in pore water chemistry (Ca2+, Mg2+, Sr, 18O and 87Sr/86Sr; Mottl et al., 1983) in

sediments associated with carbonate platforms at DSDP Site 541 (A) and Site 504 (B).

Insert shows the location of the sites.

19- Examples of marine diagenesis in platform derived sediments from Leg 166 (Fig. 10).

Scale bar = 500 m. (A) Sample of unconsolidated sediment near the top of Site 1006

(10.57 mbsf). The sediment is composed of pelagic material (foraminifera), molluscs and

calcareous mud (brown coloured material, mainly aragonite). It is uncemented and

shows no evidence of dissolution and cementation. (B) Sample from Site 1003 and a

depth of 140 mbsf, showing some dissolution and precipitation. (C) Samples from Site

1003 (483 mbsf) showing extensive infilling and dissolution. (D) Sample from Site 1003

and a depth of 601 mbsf. (E) Sample from Site 1003 (679 mbsf) showing extensive

dissolution and precipitation and a fracture which has been infilled with celestite. (F)

Sample from Site 1003 (734 mbsf) showing extensive dissolution and precipitation of

precursor carbonates as well as the precipitation of dolomite.

20- Changes in pore water chemistry (Ca2+, Mg2+, Sr, 18O and 87Sr/86Sr) in sediments

associated with carbonate platforms at ODP Site 817 (A) and Site 823 (B). The predicted

Sr concentration in diagenetic calcite is based on the Sr/Ca ratio in the pore fluids and a

distribution coefficient of 0.04. See Fig. 15 for location of sites.

21- (A) Changes in pore water chemistry (Ca2+, Mg2+, Sr, 18O and 87Sr/86Sr) in pore fluids

associated with carbonate platforms at ODP Sites 1005. The predicted Sr concentration

in diagenetic calcite is based on the Sr/Ca ratio in the pore fluids and a distribution

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 coefficient of 0.04. (B) Changes in pore water chemistry (Ca2+, Mg2+, Sr, 18O, Cl- and 13C)

in Site 1027. See Fig. 15 for location of sites.

22- Contour map of the 18Owvalues of the pore fluids from ODP Leg 166 (Swart, 2000). The

18O values of pore waters become more positive with depth as a result of alteration of

carbonate minerals at higher temperatures. See Fig. 12 for location of sites.

23- Modelled (blue line) and measured changes in the 18O values of the pore fluids

(diamonds; Swart, 2000), bulk sediments (black line and black symbols) and diagenetic

carbonate (red line) from ODP Site 1006. Model 18O data were calculated using a model

adapted from Lawrence (1989) by Stout (1985). The pore water data agrees well with

the model with the exception of the increase in the 18Owvalues, interpreted to be a

result of increases in the oceanic 18Owassociated with the last glacial period which has

diffused downward into the sediments (Schrag et al., 2002). Variations in the 18O values

of the sediments falling away from the modelled line are interpreted to be a result of

variations in the original composition of the sediments.

24- The change in the concentration of Sr (black circles) of the pore fluids (red circles) and

the 87Sr/86Sr ratios of pore fluids and dolomites (diamonds) from ODP Site 812 drilled off

the Queensland plateau (Elderfield et al., 1993; McKenzie et al., 1993). See Fig. 18 for

location of sites.

25- Concentration of Cl- in pore waters in cores drilled adjacent to a continental margin

showing the influence of saline fluids derived from the continent in the pore waters

(Feary et al., 1998). The Cl- concentrations cross cut the sedimentary layers as revealed

by the seismic data.

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26- Possible example of TSR within a mainly dolomitised oil reservoir (scale Bar = 100 m).

The sample has been stained with Alizarin red (Friedman, 1959) to show the presence of

calcite (red). Dolomite is brown. The hydrocarbons react with anhydrite (white) to

produce pyrite (black blobs) and calcite. Typically fringes of calcite are found

surrounding anhydrite. In this example the 13C value of the calcite could not be

measured and the impact upon the bulk 13C value was insignificant.

27- Relationship of the mean 13C values of the sediments in seismic sequences from ODP

Sites 1007, 1006, 1005, 1004, Clino and Unda relative to data from the same sequence in

Site 1003 (Swart & Eberli, 2005). Error bars represent the standard deviation of the data

within each sequence. The data from the ODP sites on a 1:1 line indicating the 13C

values correlate from site to site, the data from Clino and Unda fall off the line because

they have been influenced by meteoric diagenesis, rendering them useless for

chronostratigraphy.

28- The changes in the 13C values of organic and inorganic carbon from Clino (See Fig. 9 for

location). Data from Oehlert & Swart (2014). Note the strong covariation between the

two signals.

29- The potential effect of diagenesis on the 34S values of CAS in a bulk carbonate (solid

circles), assuming that BSR reduces the SO42- concentration in the interstitial pore water

(diamonds) from 28 mM (initial 34S = +22‰) to 0 mM at a depth of 1000 mbsf and that

the carbonate is completely dissolved and reprecipitated over the same interval. It is

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 assumed that the carbonate initially contained 4000 ppm SO42- and that the sediment

had a porosity of 50%.

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