the δ13c of biogenic methane in marine sediments: the influence of corg deposition rate

Ž .Chemical Geology 152 1998 139–150

The d13C of biogenic methane in marine sediments: the influence

of C deposition rateorg

Neal Blair )

Department of Marine, Earth and Atmospheric Sciences, North Carolina State UniÕersity, Raleigh, NC 27695-8208, USA

Abstract

The d13C of biogenic methane produced in marine sediments ranges from y110 to y55‰. The isotopic composition of

Ž 13 .the methane d C is constrained by the fraction of metabolized organic carbon converted to CH . The flux of labileCH 44

Ž . Ž . sorganic carbon into the seabed J and the availability of oxidants A , such as O and SO , dictate that fraction, i.e.,MOC ox 2 4Žproportionately more methane should be produced as the ratio J rA increases. Chemical, physical e.g., sedimentMOC ox

. Ž . 13resuspension and biological bioturbation and bioirrigation processes determine A . Given that d C is always lessox CH4

thand13C of the metabolizable organic carbon, d

13C should increase when J and the portion of metabolized carbonCH MOC4

Ž 2 . 13converted to methane increase. A positive linear correlation r s0.92 is observed between d C and J for aCH MOC4

Ž s .database containing four continental margin sites. When the pore water sulfate gradient DSO rDdepth is used as a4

surrogate for J , the data set is extended to 15 locations spanning all latitudes. A linear relationship between the sulfateMOC13 Ž 2 .gradient and d C r s0.98 for shelfrslope environments suggests that either J or J rA is the masterCH MOC MOC ox4

variable that controls the 13Cr12C content of the biogenic methane. Carbonate precipitation andror a methanogenic backreaction may obscure the correlation in deep-sea sediments. Evidence for the relationship between d

13C and JCH MOC4

appears to be preserved in Miocene-age dolomitic deposits. q 1998 Elsevier Science B.V. All rights reserved.

Keywords: d13C; Biogenic methane; Marine sediments

1. Introduction

The carbon isotopic composition of biogenicmethane is notable both in its extreme variability andits 13C-depletion relative to most other carbonaceousmaterials. The d

13C of marine biogenic methane, forŽexample, has a documented range of 60‰ Fig. 1;

Whiticar et al., 1986; Alperin et al., 1988; Blair et. 13al., 1994; Blair and Aller, 1995 . The C-depletion

results from kinetic isotope effects associated with

) Corresponding author. Fax: q1 919 515 7802; e-mail:[email protected]

Žthe biosynthesis of the methane Blair et al., 1993.and references within . The factors responsible for

the variability in 13Cr12 C content have not beenidentified entirely.

Microbial consortia anchored by methanogenicarchaebacteria ferment organic matter to methane viaseveral pathways. The two most common reactions,CO -reduction and acetate dissimilation, are summa-2

rized by the generalized stoichiometric relationships,

C H O q6H O™6CO q12H6 12 6 2 2 2

™3CH q3CO q6H O 1Ž .4 2 2

C H O ™3CH CO H™3CH q3CO 2Ž .6 12 6 3 2 4 2

0009-2541r98r$ - see front matter q 1998 Elsevier Science B.V. All rights reserved.Ž .PII: S0009-2541 98 00102-8

( )N. BlairrChemical Geology 152 1998 139–150140

Fig. 1. The frequency of measurement of marine biogenic d13CC H4

Ž . Ž .values. Data are from Whiticar et al. 1986 , Alperin et al. 1988 ,Ž . Ž .Blair et al. 1994 and Blair and Aller 1995 .

Methanogens rely on a limited suite of substrates,such as H and acetate, for which they are dependent2

Ž .on other fermenting microorganisms Klass, 1984 .The inability of methanogens to compete for thosesubstrates with terminal oxidizers, such as Fe, Mnand sulfate reducers, precludes most methane pro-duction in environments containing those oxidantsŽSchonheit et al., 1982; Lovley and Klug, 1983;¨Ingvorsen and Jørgensen, 1984; Robinson and Tiedje,1984; Lovley and Phillips, 1986; Chapelle and Lov-

.ley, 1992 . In both freshwater and marine sediments,methane production commences at depths in theseabed where the oxidants have been depletedŽClaypool and Kaplan, 1974; Martens and Berner,

.1974, 1977; Lovley and Klug, 1986 .Early models of methanogenesis and the associ-

ated isotopic fractionations treated the sedimentaryenvironment as a closed-system in which there wasan implied net consumption of dissolved inorganic

Ž .carbon DIC by methane production and carbonateŽ .precipitation Claypool and Kaplan, 1974 . The

closed system assumption, although reasonable forŽ .deep production of methane in the seabed )5 m ,

is untenable for shallower systems where significantsolute and gas exchange occurs with the overlying

Ž .water Martens and Klump, 1980, 1984 . Net con-sumption of DIC does not occur when organic matteris fermented to methane as indicated by the stoi-

Ž . Ž .chiometry of Eqs. 1 and 2 , laboratory culturesŽ .Tarvin and Buswell, 1934 , sediment incubationsŽ . ŽBoehme, 1993 , and field observations Boehme,

.1993; Boehme et al., 1996 . Carbonate precipitationcan be a significant removal mechanism for DIC in

Ž .some settings Claypool and Kaplan, 1974 and willbe discussed in greater detail later.

The relative contributions of CO -reduction and2Ž Ž . Ž ..acetate dissimilation Eqs. 1 and 2 to the total

production of methane appear to be significant fac-13 Žtors in controlling d C values Whiticar et al.,

.1986 . Methane produced via the acetate dissimila-tion pathway tends to be enriched in 13C relative tothat from CO -reduction in part because the fraction-2

Ž12 13 . Žation factor kr k is smaller 1.03 for acetatedissimilation vs. 1.03–1.06 for CO -reduction;2

Games et al., 1978; Fuchs et al., 1979; Belyaev etal., 1983; Balabane et al., 1987; Blair and Carter,

.1992 . In addition, the fraction of substrate convertedto methane is typically larger for acetate dissimila-tion, thus the expressed isotope effect is reduced due

Žto a mass balance constraint Blair and Carter, 1992;.Blair et al., 1993 . The dual pathway model has been

invoked to explain the isotopic differences betweenfreshwater and marine methanes, gases produced invarious freshwater environments, and seasonal varia-tions of d

13C at marine and freshwater sitesCH 4

ŽBurke and Sackett, 1986; Martens et al., 1986;Whiticar et al., 1986; Chanton and Martens, 1988;Burke et al., 1992; Kelley et al., 1992; Lansdown et

.al., 1992 . It is not clear that the range of methaned

13C values found in the marine sediments can beexplained completely by variations in the relativerates of the two pathways. Even if it could, thereremains the greater question of what controls therelative rates of those processes. In other words,what is the master variable that controls the 13Cr12 Ccomposition of marine biogenic methane?

2. The mass balance model

During early diagenesis, organic carbon is con-verted to DIC and, under the appropriate conditions,

( )N. BlairrChemical Geology 152 1998 139–150 141

Fig. 2. A simplified schematic of biogeochemical zonation in amarine sediment, omitting the depth intervals influenced by deni-trification and metal reduction. The fraction of metabolizable

Ž .organic carbon converted to methane f is dependent on howmuch of the carbon reaches the methanogenic zone. The depth ofpenetration of the fermentable carbon is controlled by its deposi-

Ž . Ž .tion rate J . Oxidant availability A dictates the thicknessMO C oxŽ .of the oxidizing zones that produce CO DIC only.2

Ž . 13 12methane in sediments Fig. 2 . The Cr C contentof the methane is constrained by a mass balancerequirement:

d13C f fd13C q 1y f d

13C 3Ž . Ž .MOC CH DIC4

where d13C , d

13C and d13C are the car-MOC CH DIC4

bon isotopic compositions relative to PDB of themetabolizable carbon, methane and DIC. The frac-tion of labile organic carbon converted to CH is4

denoted by f. The organic matter that is degraded toDIC and CH is considered to be the metabolizable4

pool. Potentially reactive material that is otherwisepreserved by absorption, geopolymerization or other

Ž .processes Hedges and Keil, 1995 is excluded fromconsideration by this definition. These alternativeprocesses, to the extent that they occur inmethanogenic marine sediments, do not influence theisotopic composition of the metabolizable fraction

Ž .appreciably Boehme, 1993; Boehme et al., 1996 .The d

13C can be described further by:CH 4

d13C fd

13C y 1y f a y1 103 4Ž . Ž . Ž .CH MOC MOC4

Ž12 13 .where a is the fractionation factor kr kMOC

associated with methanogenesis relative to d13C .MOC

This definition of a differs from the convention ofprevious studies in which the fractionation factorwas related to a specific methane precursor, such as

ŽCO or acetate Blair and Carter, 1992; Blair et al.,2.1993 . The fractionation factor associated with the

production of DIC is assumed to be 1.000.The deposition rate of metabolizable organic car-Ž .bon J is an important control of f , and byMOC

inference, d13C . The total rate of production ofCH 4

CH and DIC in marine sediments is equivalent to4ŽJ under steady state conditions Martens andMOC

.Klump, 1984 . The stoichiometries of reactions 1and 2 set an upper limit for how much CH is4

Ž .produced relative to DIC fs0.5 . However, DICŽ .production will always be favored i.e., f-0.5

because the reactive organic matter must traverse aŽ .series of oxidizing DIC-producing zones before

Ž .burial into methanogenic sediments Fig. 2 . As thedepositional flux increases, and proportionately morematerial reaches the methane-producing zone, f in-creases. The d

13C is always less than d13CCH MOC4

because a )1. Thus, d13C approachesMOC CH 4

d13C as f increases.MOC

Ž .The availability of oxidant A influences f asox

well by controlling the depth of penetration of theoxidizing zones. Chemical, physical and biologicalprocesses can influence A . Methane productionox

typically occurs deeper within the seabed relative tofreshwater sediments. The difference is a function, inpart, of the availability of sulfate in the two environ-ments. Physical resuspension of the seabed canthicken the oxidizing layer by regenerating depletedelectron-acceptors, such as Feq3 and Mnq4 , via theexposure of reduced diagenetic products to oxy-

Ž .genated water Aller et al., 1986, 1991 . Resuspen-sion thus effectively increases A . Infaunal activi-ox

ties increase the exposure of the seabed to oxidantsŽas well via irrigation and particle mixing Aller,

.1982 and thus drive the zone of methane productionŽ .to greater depths Martens, 1976 . In some situations

however, animals may decrease the exposure of or-ganic matter to oxidants by rapidly transporting ma-terial to the subsurface and thereby enhance anaero-

Žbic activities Rice, 1986; Canfield, 1989; Blair et.al., 1996 . Animal activities are supported by the

supply of food-quality organic matter. Consequently,in bioturbated sediments A can be dependent onox

J .MOC


Ž . Ž .It follows from Eqs. 3 and 4 , and the relation-ship between f and J , that a correlation betweenMOC

d13C and J should exist with the caveat thatCH MOC4

variations in a and d13C do not obscure theMOC MOC

13 Žcorrelation. Variations in d C are small ;MOC

y24"6 for typical phytoplankton across all lati-.tudes; Sackett, 1991 and Refs. within compared to

d13C and thus are not expected to be a majorCH 4

factor. This is not necessarily the case for a asMOC

the range in reported fractionation factors for CO -2

reduction, as a comparison, is large. The potentialinfluence of a on d

13C will be discussed inMOC CH 4

greater detail later. The correlation between d13CCH 4

and J was sought using previously measuredMOC

data.

3. Parameter selection and methods

3.1. Selection of d13CC H4

The carbon isotopic composition of methane typi-cally varies as a function of depth within the sedi-

Ž .ment column Fig. 3 . Much of this variation can be

Fig. 3. Depth profile of d13C values from Cape LookoutCH 4

Ž . 13Bight, NC Boehme, 1993; Boehme et al., 1996 . The C-enrich-ment at the sediment–water interface is due to aerobic methaneoxidation.

attributed to either aerobic methane oxidation nearŽthe sediment–water interface Happell et al., 1994;

.Boehme et al., 1996 or anaerobic methane oxidationŽat the base of the sulfate reduction zone Whiticar

and Faber, 1986; Alperin et al., 1988; Blair and.Aller, 1995 . Both processes selectively utilize

12 CH , and thereby enrich the residual CH pool in4 413 13 ŽC. Downcore changes in substrate d C Claypooland Kaplan, 1974; Claypool and Threlkeld, 1983;

.Galimov et al., 1983 or relative rates ofmethanogenic pathways may alter the 13CH r12 CH4 4

ratio as well.To avoid the influence of oxidation, d

13C valueswere used from the deepest portion of a core, whichin many situations was from a depth where theisotopic composition approaches an asymptotic value.

13 Ž .The asymptotic d C d is a good approximation`

Ž ."1‰ of the isotopic composition of the totalmethane pool produced over the entire sediment

Ž .column Boehme, 1993; Boehme et al., 1996 .

3.2. Estimating JM OC

The depth-integrated rate of remineralization, orŽ s.the resultant fluxes of oxidants e.g., O , SO and2 4

Ž .degradation products DIC, CH provide measures4Žof J under near steady state conditions Aller,MOC

.1980; Martens and Klump, 1984 . A combination ofthose approaches has been used to estimate J atMOC

four locations where CH d13C values were available4

Ž .Table 1 .Pore water sulfate gradients were used as surro-

gate indicators of J in an effort to extend theMOC

database to additional sites. Sulfate concentrationgradients within the seabed result from a balance ofsupply and demand. Sulfate is delivered from theoverlying water by molecular diffusion and, in someinstances, physically or biologically mediated parti-

Ž .cle mixing and solute transport Berner, 1980 . Con-sumption is via microbial sulfate reduction, which isdriven by the supply of metabolizable organic carbonŽ .Berner, 1980; Chanton and Martens, 1987 . Sulfate

Ž .reduction rates and rate constants , and the resultingsulfate gradients positively correlate with sedimentaccumulation rates from a wide variety of marine

Ženvironments Goldhaber and Kaplan, 1975; Toth


Table 1The d

13C values of methane and DIC, the flux of metabolizable organic carbon and the effective sulfate gradient from marine sediments13 13 sLocation CH d C DIC d C J mol C DSO rD z References4 MOC 4

y2 y1m yr mMrcm

Amazon Shelf y84 y8.8 18.0 0.075 Blair and Aller, 1995;Aller et al., 1996

Blake Outer Ridge y66"2 0.02"0.004 Brooks et al., 1983;Claypool and Threlkeld, 1983;Jenden and Gieskes, 1983;Galimov et al., 1983

Bransfield Strait y87"2 0.08"0.03 Whiticar and Faber, 1986Cape Hatteras, NC slope y87 4.1 0.3 Blair et al., 1994Cape Lookout Bight, NC y60"1 q9.3 45.0 2.7"0.7 Martens and Klump, 1984;

Martens et al., 1986;Boehme, 1993;Boehme et al., submitted

Cariaco Trench y67"4 0.06"0.003 Claypool et al., 1973;Lyon, 1973;Presley et al., 1973

Mississippi delta y82"2 0.45 Whelan et al., 1978Mississippi fan y75 0.005 Ishizuka et al., 1986;

Pflaum et al., 1986;Presley and Stearns, 1986

NW Africa y63"2 0.003"0.002 Barnes et al., 1979;Whelan, 1979

Orca Basin y74 0.018 Burke et al., 1986;Ishizuka et al., 1986;Pflaum et al., 1986;Presley and Stearns, 1986

y88 y24.7 0.056 Sackett et al., 1979Pigmy Basin y72 0.0038 Ishizuka et al., 1986;

Pflaum et al., 1986;Presley and Stearns, 1986

Saanich Inlet, B.C. y56 q17 3.1"1.0 Nissenbaum et al., 1972Santa Barbara Basin y85"2 0.32"0.01 Whiticar and Faber, 1986Skan Bay, AL y78"3 q3.0 16.0 0.7"0.02 Alperin et al., 1988, 1992South Guaymus q5.0 0.145 Claypool and Kaplan, 1974

The uncertainties are approximate and either represent the range in reported values or the potential error associated with estimating thevalues from irregular, or low resolution profiles.

.and Lerman, 1977; Berner, 1978; Canfield, 1989 .The correlation reflects the dependence of J onMOC

sediment deposition.Sulfate concentration profiles in the seabed range

from exponential to linear in shape. To simplify andmake consistent their parameterization, a line wasinterpolated between the points at the sediment–water

w sx Žinterface and the depth where SO ;1 mM Fig.4.4 . The slope of the line represents an effective

Ž s .whole core gradient DSO rD z . The 1-mM sulfate4

level is near the limit of detection of dissolved

sulfate for the most commonly used analytical methodŽ .BaSO precipitation . It is also the approximate4

concentration where methanogenesis becomes evi-Ž .dent Martens and Berner, 1974 . In those situations

where sulfate data were not available, but a detailedmethane concentration profile was, an apparent sul-fate gradient could be estimated by interpolating a

Žline between the surface assuming a typical seawa-.ter sulfate concentration of 32 mM and the depth of

the appearance of methane where the concentrationof sulfate was assumed to be at or below 1 mM.


4. Results and discussion

13 Ž 2 .The d C values of methane correlate r s0.92with J at the four sites for which data wereMOC

Ž .available Fig. 5 . The sites are characterized bydifferent sedimentary regimes ranging from therapidly accumulating, non-bioturbated sediments ofCape Lookout Bight, NC to the more slowly accu-mulating, extensively mixed seabed of the Amazonshelf. The sites thus experience a range in A asox

well. The positive correlation supports the hypothesisthat J influences the isotopic composition ofMOC

methane.Methane d

13C values fall into two fieldswhen related to the effective sulfate gradientŽ .Fig. 6 . One subset, which comprises data from

Ž 2nine sites, exhibits a linear correlation r s. 130.98 between d C and sulfate gradient, whichCH 4

is consistent with d13C–J relationship. The sitesMOC

are located on the shelf and continental slope, andare characterized by the shallow production of

Ž .methane -10 m in the seabed. One site, the OrcaBasin, is notable in that the result from a second set

Fig. 4. Sulfate and methane concentrations as a function of depthŽin the deltaic sediments of the Amazon River Blair and Aller,

.1995 . The vertical sulfate profile in the upper 1–2 m is due toŽ .bioirrigation Aller and Aller, 1986; Aller and Stupakoff, 1996 .

The slope of the dotted line is the effective sulfate gradient.

Fig. 5. d13C as a function of the depositional flux of metabo-CH 4

Ž .lizable organic carbon J . The least squares regression isMO C13 Ž 2 .d C s0.67 J y91.2 r s0.92 .CH MOC4

of measurements made 2 km away within the basindoes not fall along the regression line and instead is

Žassociated with the second field of data Burke et al.,.1986 .

This second field is characterized by small sulfategradients and no apparent correlation betweend

13C and DSOsrD z. Most of the sites are lo-CH 44

cated in the deep-sea, and some are euxinic basins.The gradual slopes of the sulfate profiles reflect thelimited supply of oxidizable carbon to the seabed atthese sites. Accordingly, methanogenesis occurs atconsiderable depth and is presumably slow. Thereare several possible explanations for the lack ofcorrelation between d

13C and DSOsrD z at theseCH 44

localities.Ž .Claypool and Kaplan 1974 have argued that

carbonate precipitation can influence the d13C of

methane. 13C-enriched DIC is co-produced withmethane. As the DIC is released to the pore water, itis mixed with material that was generated earlier in

Žthe overlying oxidizing zones ;y20‰; Presley.and Kaplan, 1968; Blair and Aller, 1995 . If carbon-

ate precipitates prior to, or during methanogenesis,the DIC pool is reduced in size and the DIC createdvia methanogenesis is diluted less. The resultant DICpore water pool, and any methane produced from itby CO -reduction, will be enriched in 13C relative to2


13 Ž .Fig. 6. d C as a function of effective sulfate gradient. The linear regression for the shallow-production data open circles isCH 413 Ž s . Ž 2 .d C s10.1 DSO rD z y87.1 r s0.98 .CH 44

that from non-carbonate precipitating sediments. Theobserved 13C-enrichment of the methane from thedeep-sea sites over what was expected based on thesulfate gradients, is consistent with this explanation.The precipitation of carbonate ion is indicated byDIC andror dissolved calcium concentration profilesat some sites, such as the Blake Outer Ridge and the

ŽCariaco Trench Presley et al., 1973; Claypool and.Kaplan, 1974; Claypool and Threlkeld, 1983 .

The timing of the carbonate precipitation is criti-cal. Precipitation after methanogenesis has ceasedwill not affect the d

13C of the methane. The precipi-tation effect would be most pronounced when therate of precipitation is comparable to that ofmethanogenesis. This condition is met most easilydeep within a slowly accumulating seabed whereDIC pools have had the time to accumulate butmethanogenesis is slow.

A more speculative explanation for the deep-searesults involves methanogenesis itself. Laboratoryand field observations have indicated that the methaneproduction step is reversible, i.e.,

H qCO °CH 5Ž .2 2 4

Ž .Hoehler et al., 1994 . This reversibility may lead tonet oxidation of methane under the condition of lowpH in the presence of sulfate. Under the conditions2

Žof extremely low sulfate and low pH i.e., slow2.methane formation the back reaction may become

sufficiently important that it influences the expressedisotopic fractionation. For instance, if the back reac-tion exhibits a normal kinetic isotope effect which issmaller than that associated with the forward pro-cess, the methane would still be depleted in 13C butthe difference between the d

13C values of the CH 4

and DIC would be diminished. The methane wouldbe relatively enriched in 13C as seen in the deep-seasettings.

The hypothetical relationship between d13CCH 4

and J can be tested further by considering theMOC

isotopic composition of the DIC pool. Studies indeep-sea sediments have indicated that the d

13Cgradient of the DIC near the sediment–water inter-

Žface reflects the rain rate of oxidizable carbon Mc-. 13Corkle et al., 1985 . A progressive C-depletion was

Žfound in the pore water DIC pool of surficial upper.15 cm continental slope sediments offshore of North

ŽCarolina in response to increasing J Blair et al.,MOC.1994 . Both phenomena are due to the addition of

DIC derived from the oxidation of marine organicŽ . Ž .matter ;y20‰ to seawater bicarbonate ;0‰ .

The relationship between d13C and J shouldDIC MOC

reverse when the buried DIC pool is consideredbecause 13C-enriched material is added during


Fig. 7. d13C as a function of effective sulfate gradient. TheDIC

13 Ž .asymptotic d C d of the DIC from pore water profiles was`

used when available. The line shown was drawn parallel to therelationship between d

13C and sulfate gradient and is not theCH 4

best fit regression to the DIC data.

methanogenesis. The plot of d13C vs. sulfateDIC

Ž .gradient is consistent with that prediction Fig. 7 .The data set falls about a line deliberately drawnparallel to the CH d

13C–DSOsrD z relationship.4 4

Possible evidence for the dependency of d13CCH 4

and d13C on J has been preserved in theDIC MOC

geologic record. Quantitative paleo-indicators ofJ are virtually nonexistent. However, given thatMOC

the depositional flux is defined by,

w xJ s C v 6Ž .´MOC MOC

w xwhere C is the concentration of the metaboliz-MOC

able organic carbon in the deposited particles, and v́

is the depositional rate of sediment, then v can be a´w xrelative indicator of J through time if CMOC MOC

remains nearly invariant. A positive correlation isobserved between methane d

13C values and apparentŽ .sediment accumulation rate v in Pliocene–Miocene

age dolomites from offshore Baja and southern Cali-

Ž .fornia Fig. 8; Pisciotto and Mahoney, 1981 . Therequisite parallel relationship between carbonate d

13Cand v also exists. If the correlation is due to adependence on J , it indicates that three condi-MOC

Ž .tions were met, i.e., 1 v is roughly proportional toŽ . w xv 2 C did not change dramatically over the´ MOC

Ž .time scale of the sediment record ;14 MY , andŽ .3 the precipitation of the carbonates occurred pri-marily after methanogenesis ceased.

The correlation between d13C and sulfate gra-CH 4

dient in the continental margin sediments has impli-cations that concern other potential controls of the13Cr 12 C content of methane. The site locations spannearly all latitudes and therefore experience a widerange of temperatures. The correlation betweend

13C and DSOsrD z indicates that temperatureCH 44

is not an independent variable that influences thefractionation factors associated with themethanogenic reactions. The evidence for a tempera-ture effect from previous studies has been conflict-

Fig. 8. The d13C of methane and carbonate from Pliocene–

Miocene age dolomites as a function of the apparent accumulationŽ .rate Pisciotto and Mahoney, 1981 . Data from samples that have

experienced temperatures 0408C were not included to avoidpotential thermal effects.


ing. Whereas a potential temperature effect was sug-Žgested by at least one study Games and Hayes,

.1976; Games et al., 1978 , other investigations havefailed to document a relationship between tempera-

Ž .ture and fractionation factor Belyaev et al., 1983 .Results from a modeling study of methane produc-tion at Cape Lookout Bight have suggested that a

for CO -reduction was dependent on temperature2Ž .Blair et al., 1993 . Subsequent refinements of themodel indicate that a may be dependent on CO -re-2

duction rate, which is partially dependent on temper-Ž .ature Blair, unpublished results .

As noted previously, the fractionation factor asso-ciated with methanogenesis is large and has thepotential to be extremely variable. The fractionation

Ž .factors reported for CO -reduction 1.03–1.06 are2

from freshwater culture studies and their relevance tonatural marine settings is uncertain. Apparent frac-

Žtionation factors the isotopic separation between.coexisting CO and CH estimated from field data2 4

range from 1.05 to 1.09 in marine sedimentsŽ .Whiticar et al., 1986 . It is not known whether theactual fractionation factors would be as variablewhen corrected for the effects of transport, opensystem losses of DIC and CH , and other4

Ž .methanogenic pathways e.g., acetate dissimilation .The relationship betweend

13C and sulfate gradi-CH 4

ent indicates that either a is nearly constant inMOC

the environments represented, or it is a function ofJ rA . If a is dependent on methanogenicMOC ox MOC

Ž .rate Blair, unpublished results , it would be indi-rectly related to f , and therefore J rA .MOC ox

The isotopic composition of CH is dependent on4

the relative rates of the two methanogenic processes,CO -reduction and acetate dissimilation. The2

d13C –DSOsrD z relationship indicates that ei-CH 44

ther the relative rates vary little amongst the differentlocations, or they are dependent on J rA . Ma-MOC ox

rine methanogenesis is thought to be dominated byCO -reduction, though that belief has not been tested2

Žextensively Nissenbaum et al., 1972; Claypool andKaplan, 1974; Crill and Martens, 1986; Kuivila et

.al., 1990 . In the well-characterized sediments ofCape Lookout Bight, CO -reduction is responsible2

Žfor )85% of the methane produced annually Crill. 13and Martens, 1986 . Some of the most C-enriched

gas is produced at Cape Lookout Bight, thus basedon the hypothesized relationship between d

13CCH 4

Fig. 9. The potential controls of methane carbon isotopic composi-tion. J r A and temperature control the methanogenic rate.MO C ox

The methane production rate controls f, which in turn influencesd

13C . The methanogenic rate may also regulate the fractiona-CH 4

tion factor and the relative rates of acetate dissimilation andCO -reduction.2

Žand pathway acetate dissimilation produces a more13C-enriched gas; Whiticar et al., 1986; Blair and

.Carter, 1992 , it appears that methanogenesis is dom-inated by CO -reduction in the marine environments2

considered in this study. However, that does noteliminate the possibility that the relative rates aredependent on some function of J and A .MOC ox

Other potential interactions between the various pa-13 Ž .rameters may also influence d C values Fig. 9 .

5. Conclusions

The isotopic composition of the diagenetic prod-ucts, CH and DIC, are constrained by mass balance4

requirements. The total inventory of 13C and 12 C inthe CH and DIC pools must be equivalent to that4

delivered to the seabed as metabolizable organiccarbon. The d

13C value of each individual product isa function of its relative rate of production, which iscontrolled by the delivery rate of labile carbon to theseabed and the availability or supply of oxidants.The d

13C of both CH and DIC increase with the4

deposition rate of reactive marine organic matter intypical continental margin sediments. The relation-ship between d

13C and J may be obscured byMOC

carbonate precipitation, or more speculatively,methanogenic back reactions in situations wheremethane production occurs deep within the seabed


and is apparently very slow. These situations aremost likely to be encountered in deep-sea settings.

The mass balance constraint, and the resultingarguments likely apply to freshwater systems as well.Plant exudates and the transport of O into soils2

through rhizomes in vegetated environments, as wellas poorly defined hydrological conditions, will com-plicate the determination of J , A and d

13CMOC ox CH 4

Ž .Holzapfel-Pschorn et al., 1985; Kelley et al., 1995 .Initial studies testing the relationship between thoseparameters should be attempted in the sediments oflarge lakes to minimize those effects. Freshwatermethane is typically enriched in 13C relative to ma-rine gas and this difference has been attributed to thepresumed dominance of the acetate dissimilationpathway in the freshwater environments and CO -re-2

Ž .duction in marine sediments Whiticar et al., 1986 .However, the 13C-enrichment is also predicted by themass balance model because J may be greaterMOC

due to the proximity of vegetation, and A with lessox

sulfate is lower in freshwater settings. Clearly, addi-tional studies are needed to resolve those issues.

Acknowledgements

Many thanks to Chris Martens, Jeff Chanton,Marc Alperin, Dan Albert, Cheryl Kelley, SueBoehme, Bob Aller and many others for their helpfuldiscussions, field assistance and data, without whichthis study would not have been possible. Specialthanks go to George Claypool for pointing out theimportance of carbonate precipitation to this prob-lem. This study was supported by grants from NASAŽ . ŽNAGW-838 , NSF OCE-8812907, OCE-9115709,

.OCE-9301793 and the NOAA National UnderseaResearch Program at Wilmington, NC.

References

Aller, R.C., 1980. Diagenetic processes near the sediment–waterinterface of Long Island Sound: I. Decomposition and nutrient

Ž .element geochemistry S,N,P . Adv. Geophys. 22, 237–349.Aller, R.C., 1982. The effects of macrobenthos on chemical

properties of marine sediment and overlying water. In: Mc-Ž .Call, P.L., Tevesz, M.J.S. Eds. , Animal Sediment Relations.

Plenum, New York, pp. 53–102.

Aller, J.Y., Aller, R.C., 1986. General characteristics of benthicfaunas on the Amazon inner continental shelf with comparisonto the shelf off the Changjiang River, East China Sea. Cont.Shelf Res. 6, 291–310.

Aller, J.Y., Stupakoff, I., 1996. The distribution and seasonalcharacteristics of benthic communities on the Amazon shelf asindicators of physical processes. Cont. Shelf Res. 16, 717–752.

Aller, R.C., Mackin, J.E., Cox, R.T., 1986. Diagenesis of Fe andS in Amazon inner shelf muds: apparent dominance of Fereduction and implications for the genesis of ironstones. Cont.Shelf Res. 6, 263–289.

Aller, R.C., Aller, J.Y., Blair, N.E., Mackin, J.E., Rude, P.D.,Stupakoff, I., Patchineelam, S., Boehme, S.E., Knopper, B.,1991. Biogeochemical Processes in Amazon Shelf sediments.Oceanography 4, 27–32.

Aller, R.C., Blair, N.E., Xia, Q., Rude, P.D., 1996. Remineraliza-tion rates, recycling, and storage of carbon in Amazon shelfsediments. Cont. Shelf Res. 16, 753–786.

Alperin, M.J., Reeburgh, W.S., Whiticar, M.J., 1988. Carbon andhydrogen isotope fractionation resulting from anaerobicmethane oxidation. Global Biogeochem. Cycles 2, 279–288.

Alperin, M.J., Reeburgh, W.S., Devol, A.H., 1992. Organic car-bon remineralization and preservation in sediments of Skan

Ž .Bay, Alaska. In: Whelan, J., Farrington, J.W. Eds. , OrganicMatter: Productivity, Accumulation, and Preservation in Re-cent and Ancient Sediments. Columbia Press, New York, pp.99–122.

Balabane, M., Galimov, E., Hermann, M., Letolle, E., 1987.Hydrogen and carbon isotope fractionation during experimen-tal production of bacterial methane. Org. Geochem. 11, 115–119.

Barnes, R.O., Gieskes, J.M., Horvath, J., Akiyama, W., 1979.Interstitial water studies, Legs 47A,B. Init. Reports DSDP 47,577–581.

Belyaev, S.S., Wolkin, R., Kenealy, W.R., DeNiro, M.J., Epstein,S., Zeikus, J.G., 1983. Methanogenic bacteria from theBondyuzhskoe oil field: general characterization and analysisof stable-carbon isotopic fractionation. Appl. Environ. Micro-biol. 45, 691–697.

Berner, R.A., 1978. Sulfate reduction and the rate of deposition ofmarine sediments. Earth Planet. Sci. Lett. 37, 492–498.

Berner, R.A., 1980. Early Diagenesis, A Theoretical Approach.Princeton Univ. Press, NJ, 241 pp.

Blair, N.E., Aller, R.C., 1995. Anaerobic methane oxidation onthe Amazon shelf. Geochim. Cosmochim. Acta 59, 3707–3715.

Blair, N.E., Carter, W.D. Jr., 1992. The carbon isotope biogeo-chemistry of acetate from a methanogenic marine sediment.Geochim. Cosmochim. Acta 56, 1247–1258.

Blair, N.E., Boehme, S.E., Carter, W.D. Jr., 1993. The carbonisotope biogeochemistry of methane production in anoxic sedi-

Ž .ments: 1. Field observations. In: Oremland, R.S. Ed. , Bio-geochemistry of Global Change—Radiatively Active TraceGases. Chapman & Hall, New York, pp. 574–593.

Blair, N.E., Levin, L.A., DeMaster, D.J., Plaia, G., 1996. Theshort-term fate of fresh algal carbon in continental slopesediments. Limnol. Oceanogr. 41, 1208–1219.

Blair, N.E., Plaia, G.R., Boehme, S.E., DeMaster, D.J., Levin,


L.A., 1994. The remineralization of organic carbon on theNorth Carolina continental slope. Deep-Sea Res. II 41, 755–766.

Boehme, S.E., 1993. The carbon isotope biogeochemistry of amethanogenic marine sediment. PhD thesis, North CarolinaState University, 144 pp.

Boehme, S.E., Blair, N.E., Chanton, J.P., Martens, C.S., 1996. Amass balance of 13C and 12 C in an organic-rich methane-pro-ducing marine sediment. Geochim. Cosmochim. Acta 60,3835–3848.

Brooks, J.M., Barnard, L.A., Wiesenburg, D.A., Kennicutt, M.C.II, Kvenvolden, K.A., 1983. Molecular and isotopic composi-tions of hydrocarbons at Site 533, Deep Sea Drilling ProjectLeg 76. Init. Reports DSDP 76, 377–389.

Burke, R.A. Jr., Sackett, W.M., 1986. Stable hydrogen and carbonisotopic compositions of biogenic methanes from several shal-

Ž .low aquatic environments. In: Sohn M.L. Ed. , Organic Ma-rine Chemistry. American Chemical Society Series 305, pp.297–313.

Burke, R.A. Jr., Sackett, W.M., Brooks, J.M., 1986. Hydrogen-and carbon-isotope compositions of methane from Deep SeaDrilling Project Site 618, Orca Basin. Init. Reports DSDP 96,777–780.

Burke, R.A., Barber, T.R., Sackett, W.M., 1992. Seasonal varia-tions of stable hydrogen and carbon isotope ratios of methanein subtropical freshwater sediments. Global Biogeochem. Cy-cles 6, 125–138.

Canfield, D.E., 1989. Sulfate reduction and oxic respiration inmarine sediments: implications for organic carbon preserva-tion in euxinic environments. Deep-Sea Res. 36, 121–138.

Chanton, J.P., Martens, C.S., 1987. Biogeochemical cycling in anorganic-rich coastal marine basin: 7. Sulfur mass balance,oxygen uptake and sulfide retention. Geochim. Cosmochim.Acta 51, 1187–1199.

Chanton, J.P., Martens, C.S., 1988. Seasonal variations in ebulli-tive flux and carbon isotopic composition of methane in a tidalfreshwater estuary. Global Biogeochem. Cycles 2, 289–298.

Chapelle, F.H., Lovley, D.R., 1992. Competitive exclusion ofŽ .sulfate reduction by Fe III -reducing bacteria: a mechanism

for producing discrete zones of high-iron groundwater. GroundWater 30, 29–36.

Claypool, G.E., Kaplan, I.R., 1974. The origin and distribution ofŽ .methane in marine sediments. In: Kaplan, I.R. Ed. , Natural

Gases in Marine Sediments. Plenum, pp. 99–139.Claypool, G.E., Threlkeld, C.N., 1983. Anoxic diagenesis and

methane generation in sediments of the Blake Outer Ridge,Deep Sea Drilling Project Site 533, Leg 76. Init. ReportsDSDP 76, 391–402.

Claypool, G.E., Presley, B.J., Kaplan, I.R., 1973. Gas analyses insediment samples from legs 10, 11, 13, 14, 15, 18 and 19. Init.Reports DSDP XIX, 879–884.

Crill, P.M., Martens, C.S., 1986. Methane production from bicar-bonate and acetate in an anoxic marine sediment. Geochim.Cosmochim. Acta 50, 2089–2097.

Fuchs, G., Thauer, R., Zeigler, H., Stichler, W., 1979. Carbonisotope fractionation by Methanobacterium thermoautotroph-icum. Arch. Microbiol. 120, 135–139.

Galimov, E.M., Vernadsky, V.I., Kvenvolden, K.A., 1983. Con-centrations and carbon isotopic compositions of CH and CO4 2

in gas from sediments of the Blake Outer Ridge, Deep SeaDrilling Project Leg 76. Init. Reports DSDP 76, 403–407.

Games, L.M., Hayes, J.M., 1976. On the mechanisms of CO and2

CH production in natural anaerobic environments. In: Nriagu,4Ž .J.O. Ed. , Proc. 2nd Int. Symp. Environ. Biogeochem. Michi-

gan Science Press, Ann Arbor, pp. 51–73.Games, L.M., Hayes, J.M., Gunsulas, R.P., 1978. Methane-pro-

ducing bacteria: natural fractionations of the stable carbonisotopes. Geochim. Cosmochim. Acta 42, 1295–1297.

Goldhaber, M.B., Kaplan, I.R., 1975. Controls and consequencesof sulfate reduction rates in recent marine sediments. Soil Sci.119, 42–55.

Happell, J.D., Chanton, J.P., Showers, W.S., 1994. The influenceof methane oxidation on the stable isotopic composition ofmethane emitted from Florida swamp forests. Geochim. Cos-mochim. Acta 58, 4377–4388.

Hedges, J.I., Keil, R.G., 1995. Sedimentary organic matter preser-vation: an assessment and speculative synthesis. Mar. Chem.49, 81–115.

Hoehler, T.M., Alperin, M.J., Albert, D.B., Martens, C.S., 1994.Field and laboratory studies of methane oxidation in an anoxicmarine sediment: evidence for a methanogen-sulfate reducerconsortium. Global Biogeochem. Cycles 8, 451–463.

Holzapfel-Pschorn, A., Conrad, R., Seiler, W., 1985. Production,oxidation and emission of methane in rice paddies. FEMSMicrobiol. Ecol. 31, 343–351.

Ingvorsen, K., Jørgensen, B.B., 1984. Kinetics of sulfate uptakeby freshwater and marine species of DesulfoÕibrio. Arch.Microbiol. 139, 61–66.

Ishizuka, T., Kawahata, H., Aoki, S., 1986. Interstitial watergeochemistry and clay mineralogy of the Mississippi Fan andOrca and Pigmy Basins, Deep Sea Drilling Project Leg 96.Init. Reports 96, 711–728.

Jenden, P.D., Gieskes, J.M., 1983. Chemical and isotopic compo-sition of interstitial water from Deep Sea Drilling Project Sites533 and 534. Init. Reports DSDP 76, 453–461.

Kelley, C.A., Dise, N.B., Martens, C.S., 1992. Temporal varia-tions in the stable carbon isotopic composition of methaneemitted from Minnesota peatlands. Global Biogeochem. Cy-cles 6, 263–269.

Kelley, C.A., Martens, C.S., Ussler, W. III, 1995. Methane dy-namics across a tidally flooded riverbank margin. Limnol.Oceanogr. 40, 1112–1129.

Klass, D.L., 1984. Methane from anaerobic fermentation. Science233, 1021–1028.

Kuivila, K.M., Murray, J.W., Devol, A.H., 1990. Methane produc-tion in the sulfate-depleted sediments of two marine basins.Geochim. Cosmochim. Acta 54, 403–411.

Lansdown, J.M., Quay, P.D., King, S.L., 1992. CH production4

via CO reduction in a temperature in a temperate bog: a2

source of 13C-depleted CH . Geochim. Cosmochim. Acta 56,4

3493–3503.Lovley, D.R., Klug, M.J., 1983. Sulfate reducers can outcompete

methanogens at freshwater sulfate concentrations. Appl. Envi-ron. Microbiol. 45, 187–192.


Lovley, D.R., Klug, M.J., 1986. Model for the distribution ofsulfate reduction and methanogenesis in freshwater sediments.Geochim. Cosmochim. Acta 50, 11–18.

Lovley, D.R., Phillips, E.J.P., 1986. Organic matter mineralizationwith reduction of ferric iron in anaerobic sediments. Appl.Environ. Microbiol. 51, 683–689.

Lyon, G.L., 1973. Interstitial water studies, leg 15-Chemical andisotopic composition of gases from Cariaco Trench sediments.Init. Reports DSDP 20, 773–774.

Martens, C.S., 1976. Control of methane sediment–water bubbletransport by macrofaunal irrigation in Cape Lookout Bight,North Carolina. Science 192, 998–1000.

Martens, C.S., Berner, R.A., 1974. Methane production in theinterstitial waters of sulfate-depleted marine sediments. Sci-ence 185, 1167–1169.

Martens, C.S., Berner, R.A., 1977. Interstitial water chemistry ofanoxic Long Island Sound sediments: 1. Dissolved gases.Limnol. Oceanogr. 22, 10–25.

Martens, C.S., Klump, J.V., 1980. Biogeochemical cycling in anorganic-rich coastal basin: I. Methane sediment–water ex-change processes. Geochim. Cosmochim. Acta 44, 471–490.

Martens, C.S., Klump, J.V., 1984. Biogeochemical cycling in anorganic-rich coastal marine basin: 4. An organic carbon budgetfor sediments dominated by sulfate reduction and methanogen-esis. Geochim. Cosmochim. Acta 48, 1987–2004.

Martens, C.S., Blair, N.E., Green, C.D., Des Marais, D.J., 1986.Seasonal variations in the stable carbon isotopic signature ofbiogenic methane in a coastal sediment. Science 233, 1300–1303.

McCorkle, D.C., Emerson, S.R., Quay, P.D., 1985. Stable carbonisotopes in marine porewaters. Earth Planet. Sci. Lett. 74,13–36.

Nissenbaum, A., Presley, B.J., Kaplan, I.R., 1972. Early diagene-sis in a reducing fjord, Saanich Inlet, British Columbia: I.Chemical and isotopic changes in major components of inter-stitial water. Geochim. Cosmochim. Acta 36, 1007–1027.

Pflaum, R.C., Brooks, J.M., Cox, H.B., Kennicutt, M.C. II, Sheu,D.-D., 1986. Molecular and isotopic analyses of core gasesand gas hydrates, Deep Dea Drilling Project Leg 96. Init.Reports DSDP 96, 781–784.

Pisciotto, K.A., Mahoney, J.J., 1981. Isotopic survey of diageneticcarbonates, Deep Sea Drilling Project Leg 63. Init. ReportsDSDP 63, 595–609.

Presley, B.J., Kaplan, I.R., 1968. Changes in dissolved sulfate,calcium and carbonate from interstitial water of near-shoresediments. Geochim. Cosmochim. Acta 32, 1037–1048.

Presley, B.J., Stearns, S., 1986. Interstitial water chemistry, DeepSea Drilling Project, Leg 96. Init. Reports DSDP 96, 697–709.

Presley, B.J., Culp, J., Petrowski, C., Kaplan, I.R., 1973. Intersti-tial water studies, leg 15-major ions Br, Mn, NH , Li, B, Si,3

and d13C. Init. Reports DSDP 20, 805–809.

Rice, D.L., 1986. Early diagenesis in bioadvective sediments:relationships between the diagenesis of beryllium-7, sedimentreworking rates, and the abundance of conveyor-belt depositfeeders. J. Mar. Res. 44, 149–184.

Robinson, J.A., Tiedje, J.M., 1984. Competition between sulfate-reducing and methanogenic bacteria for H under resting and2

growing conditions. Arch. Microbiol. 137, 26–32.Sackett, W.M., 1991. A history of the d

13C composition ofoceanic plankton. Mar. Chem. 34, 153–156.

Sackett, W.M., Brooks, J.M., Bernard, B.B., Schwab, C.R., Chung,H., Parker, R.A., 1979. A carbon inventory for Orca Basinbrines and sediments. Earth Planet. Sci. Lett. 44, 73–81.

Schonheit, P., Kristjansson, J.K., Thauer, R.K., 1982. Kinetic¨mechanism for the ability of sulfate reducers to outcompetemethanogens for acetate. Arch. Microbiol. 132, 285–288.

Tarvin, D., Buswell, A.M., 1934. The methane fermentation oforganic acids and carbohydrates. J. Am. Chem. Soc. 56,1751–1755.

Toth, D.J., Lerman, A., 1977. Organic matter reactivity andsedimentation rates in the ocean. Am. J. Sci. 277, 465–485.

Whelan, J.K., 1979. C to C hydrocarbons from IPOD holes 3971 7

and 397A. Init. Reports DSDP 47, 531–539.Whelan, T., III, Bernard, B.B., Brooks, J.M., 1978. Carbon iso-

tope variations in total carbon dioxide and methane frominterstitial waters of nearshore sediments. In: Robinson, B.W.Ž .Ed. , Stable Isotopes in the Earth Sciences. New ZealandDepartment of Scientific and Industrial Research Bulletin 220,Wellington, NZ, pp. 39–47.

Whiticar, M.J., Faber, E., 1986. Methane oxidation in sedimentand water column environments-isotopic evidence. Org.Geochem. 10, 759–768.

Whiticar, M.J., Faber, E., Schoell, M., 1986. Biogenic methaneformation in marine and freshwater environments: CO reduc-2

tion vs. acetate fermentation– Isotope evidence. Geochim.Cosmochim. Acta 50, 693–709.

the δ13c of biogenic methane in marine sediments: the influence of corg deposition rate

Documents