rates of organic carbon oxidation in deep sea sediments in the eastern north atlantic from pore...

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Marine Geology 212

Rates of organic carbon oxidation in deep sea sediments

in the eastern North Atlantic from pore water profiles

of O2 and the d13C of dissolved inorganic carbon

S. Papadimitriou*, H. Kennedy, D.N. Thomas

University of Wales-Bangor, School of Ocean Sciences, Menai Bridge, Anglesey LL59 5AB, UK

Received 18 December 2003; received in revised form 17 May 2004; accepted 26 August 2004

Abstract

The remineralization rate of sedimentary organic carbon (Rorg) and the depth-integrated, diffusion-supplied O2 consumption

rate (IOC) during microbial metabolism in sediments was investigated in three deep-sea sites at 1100, 2000 and 3500 m water

depth in the eastern north Atlantic during the spring and summer 1998. In-situ pore water O2 profiles yielded an IOC of

0.45F0.07 mmol O2 m�2 day�1 at the deepest site (n=3) and ca. 1–1.5 mmol O2 m

�2 day�1 at the shallowest site (n=2). The

Rorg was independently estimated at all three sites from ex-situ pore water profiles of the isotopic composition of total dissolved

inorganic carbon (d13CT), assuming that the concentration and isotopic composition of pore water CTwith depth in the sediment

was controlled only by microbial oxidation of isotopically depleted sedimentary organic carbon. The Rorg was thus estimated to

be ca. 0.5–0.6 mmol C m�2 day�1 at the shallowest site and ca. 0.3–0.4 mmol C m�2 day�1 at the two deeper sites.

Stoichiometric and isotopic constraints indicated that oxic remineralization of sedimentary organic matter was the dominant

metabolic pathway in the sediments at 3500 m water depth. Similarly, stoichiometric and isotopic constraints suggested that the

Rorg estimates from the ex-situ pore water d13CT profiles from 1100 and 2000 m water depth were likely to be minimum values

and provided evidence for the occurrence of post-oxic remineralization processes. Post-oxic metabolism in the sediments of

these sites could be linked to, or even augmented by, the non-diffusive mode of supply of organic matter mediated by infaunal

organisms below the oxic sediment layer.

D 2004 Elsevier B.V. All rights reserved.

Keywords: marine sediments; carbon cycling; carbon isotopes; oxygen; North Atlantic

0025-3227/$ - see front matter D 2004 Elsevier B.V. All rights reserved.

doi:10.1016/j.margeo.2004.08.003

* Corresponding author. Tel.: +44 0 1248 38 8116.

E-mail address: [email protected]

(S. Papadimitriou).

1. Introduction

Chemical transformations in modern marine sedi-

ments are directly or indirectly induced by microbial

metabolism. Its dominant component is driven by the

microbial oxidation of the particulate organic matter

(2004) 97–111

S. Papadimitriou et al. / Marine Geology 212 (2004) 97–11198

(POM) that rains onto the sea floor following

biological production in the euphotic zone of the

ocean and subsequent incorporation in the sediments

by sedimentation and mixing. The microbial oxidation

of POM proceeds through sequential utilization of

oxidants in the order of the most energy-yielding in

more or less discrete redox reaction zones in sedi-

ments (Berner, 1980). In the deep-sea environment,

the majority of the oxidation occurs within the upper

few centimetres mostly by utilization of dissolved

molecular oxygen (O2) (e.g., Emerson and Bender,

1981; McCorkle and Emerson, 1988). The amount

and reactivity of POM that escapes oxidation in the

oxic zone of deep-sea sediments often supports

comparatively little activity in deeper sediment

sections via progressively less energy-yielding oxi-

dants, such as nitrate (NO3�), oxides of manganese

and iron, and sulphate (SO42�) (Froelich et al., 1979).

The seasonal enhancement of the supply and

quality of sedimentary POM is associated with

phytoplankton blooms in the euphotic zone and is

manifest in deposition of phytodetrital material (Beau-

lieu, 2002, and references therein). Microbiological

studies on phytodetrital material (Lochte and Turley,

1988; Smith et al., 1996) demonstrated the potential

for coupling of seasonal patterns in deposition to

benthic metabolism. This was supported further by

measurements of sediment O2 consumption, with

enhanced rates within the period of development

and sinking of the surface plankton bloom in the

eastern north Atlantic (Patching et al., 1986; Pfann-

kuche, 1993). Similar measurements in the same and

other parts of the World Ocean did not detect a prompt

response to seasonal pulses of POM onto the seafloor

from metabolism in sediments (Sayles et al., 1994;

Lampitt et al., 1995; Smith et al., 1998). In those

cases, the decoupling was demonstrated to be due to a

number of interacting factors, including the low

reactivity of the deposited POM and heavy utilization

of phytodetrital aggregates by epibenthic fauna before

incorporation into underlying sediments.

During sediment metabolism, a fraction of the

carbon and nitrogen incorporated in the POM during

biological production in the euphotic zone of the

oceans is released to the sediment pore water in the

form of their dissolved inorganic species. The carbon

incorporated in the POM (POC) in the surface ocean

water is depleted in the heavy isotope, 13C, with values

of d13CPOC varying latitudinally between �35x and

�18x (e.g., Hofmann et al., 2000). Consequent on the

metabolic POC oxidation is an increase in the

concentration of total dissolved inorganic carbon

(CT) in the pore water with depth in the sediment.

This is accompanied by a progressively 13C-depleted

stable carbon isotope ratio of pore water CT (d13CT)

relative to that in the bottom ocean water overlying the

sediments (d13CT, BW). The extent of the modification

of pore water d13CT towards isotopically depleted

values will depend on the oxidation rate of POC. The

occurrence and contribution to the pore water CT pool

of dissolution of carbonate minerals (CaCO3) in the

oxic zone of sediments will moderate this modifica-

tion, because sedimentary carbonate carbon is consid-

erably more enriched isotopically than POC (i.e.,

d13CCaCO3~0–1x) (McCorkle and Emerson, 1988;

Sayles and Curry, 1988; Martin et al., 2000). Hence,

measurement of the stable isotope ratio of different

carbon pools has become a powerful tool in the study

of metabolic processes in sediments.

The present study was conducted in the deep (N1000

m water depth) eastern north Atlantic during the UK

community thematic programme BENBO (Biogeo-

chemistry in the Deep Benthic Boundary Layer). As

part of the programme, the oxidation of POC during

sediment metabolism was investigated at three sites

during the period of development and sinking of the

spring phytoplankton bloom. To this end, in-situ

profiles of pore water O2 (Black et al., 2001) were

fitted to a steady state reaction-diffusion equation to

determine the depth-integrated O2 consumption in the

sediments. The rate of POC oxidation was quantified

by applying isotopic mass balance to ex-situ measure-

ments of pore water d13CT. The calculated rates are

qualitatively assessed by comparison with past meas-

urements of sediment metabolism in the area, while the

occurrence of CaCO3 dissolution near the sediment

surface and the indirect role of benthic infauna in the

oxidation process are also discussed.

2. Sampling and analytical methods

2.1. Study area

Three sites located between 528 and 578N in the

eastern north Atlantic (Fig. 1) were visited during

Fig. 1. Study sites in the eastern north Atlantic. Filled diamonds indicate the sites that were sampled for sediments during the BENBO cruises in

spring–summer 1998. Open circles indicate the water column stations sampled during the METEOR cruise in May 1997 (Prof. A. Koertzinger,

University of Kiel, Germany).

S. Papadimitriou et al. / Marine Geology 212 (2004) 97–111 99

cruise 111 in spring 1998 and cruise 113 in summer

1998 on RSS Charles Darwin. The shallowest site B

(1100 m) is located on the west flank of the Rockall

Plateau near the Hatton-Rockall Basin, the intermedi-

ate site C (1965 m) is on the east flank of the Rockall

Plateau and the deepest site A (3560 m) is in the

mouth of Rockall Trough north of the Porcupine

Abyssal Plain. Sediment lithology, deposition and

hydrography at the sites were discussed elsewhere

(Thomson et al., 2000). Briefly, sediment deposition

at the sites is slow, with a sedimentation rate of 4.4,

6.5 and 2.1 cm ky�1 at sites B, C and A respectively.

All sites have a deep surface mixed layer, and a

biodiffusive particle mixing rate of 0.088 and 0.045

cm2 year�1 was quantified from 210Pbexcess profiles at

sites B and A, respectively. At sites B and C, the210Pbexcess and fallout radionuclide profiles indicated a

non-diffusive particle mixing event, which was

attributed to echiurian or sipunculid worms retrieved

from these sediments. Openings of animal burrows of

various sizes were observed in seabed photographs,

and segments of deep-penetrating (down to 19 cm

below the sediment–water interface) burrows were

recovered in box cores at sites B and C, containing

faecal pellets and also a green-coloured slurry at site

B. Infaunal abundance in the upper 10 cm of the

sediment was highest site B and lowest at site A (D.

Hughes, unpublished data).

Based on their carbonate content in the upper 40

cm, the sediments at site B are carbonate oozes (80%

CaCO3), while those at site C are carbonate marls

(54% CaCO3) (Thomson et al., 2000). At the deepest

site, 12–15 cm of carbonate ooze (76% CaCO3)

overlies glacial clays (20% CaCO3) (op. cit.). At all

sites, the sediments were low in organic matter, with

core top maximum concentrations of 0.56% and

0.65% at site B, 0.61% and 0.71% at site C and

0.31% and 0.36% at site A (Thomson et al., 2000;

Papadimitriou et al., 2002). The porosity (/) was

invariable over the upper 36–40 cm of the sediment at

0.68F0.05 (1r, n=58), 0.70F0.03 (n=55) and

0.70F0.04 (n=61) at sites A, B and C respectively

(from wet and dry bulk sediment densities in

Thomson et al., 2001).


2.2. Analytical methods

Sediment sampling took place twice on each cruise

at all three sites (Table 1). In-situ O2 profiles were

measured with a vertical resolution of 0.025 cm using

microelectrodes attached to a benthic lander (Black et

al., 2001). All sediment cores were retrieved with a

multicorer (Barnett et al., 1984), using 11 cm diameter

core tubes. Immediately after recovery, sediments

were sectioned every 0.5 cm down to 3 cm below the

sediment–water interface, every 1 cm below this depth

and down to 10 cm depth in the sediment, and every 2

cm thereafter. The pore waters were separated by

centrifugation at in-situ temperature and under nitro-

gen atmosphere. Sample collection for the determi-

nation of pore water d13CT and its analysis followed

the procedure described in Papadimitriou et al. (2004).

The sediment solids were kept frozen for subsequent

analyses onshore. The isotopic composition of bulk

sedimentary POC (d13CPOC) was determined in sedi-

ment sections from the upper 11 cm of two cores from

site C and from the upper 35 cm of a single core from

site A as described in Kennedy et al. (2004). The

d13CPOC and d13CT are reported relative to Vienna Pee

Dee Belemnite (VPDB) as d13Csample=1000[(Rsample/

RVPDB)�1], where R=13C/12C, and were measured on

a EUROPA-PDZ 20/20 mass spectrometer, with a

Table 1

Site details and date of collection of cores processed and analyzed for po

Site Depth (m) ha (8C) Station

B 1104 5.4 54407#3

54702#8

1101 #010b

211c 1085 5.8

C 1965 3.2 54401#10d

54409#2

54701#5d

207c 1874 3.6

A 3560 2.4 54403#2d

54703#1

3573 #004b

3357 #012b

205c 2784 2.8

The details of the deployment location and date of O2 micro-profiles were

water depth.a Bottom water temperature.b O2 microelectrode profiles.c Provided by Prof. A. Koertzinger, University of Kiel, Germany.d Analyzed for d13CPOC.

precision of F0.1x based on internal carbonate

standards.

3. Calculation of oxidation rates of sedimentary

organic carbon

3.1. Pore water O2

The depth-integrated O2 consumption rate (IOC) in

the sediments was estimated by fitting a steady state

diffusion-reaction equation to the pore water O2

profiles (e.g., Rasmussen and Jorgensen, 1992) for

constant porosity (see Section 2.1). At the slow

sedimentation and particle mixing rates of the deep

sea, assuming that bio-irrigation is insignificant,

transport by molecular diffusion and reaction primar-

ily control the pore water concentration of dissolved

oxidants consumed and metabolites produced during

the bacterial oxidation of sedimentary organic matter.

In the model, the O2 consumption rate declines

exponentially with depth in the sediment and is

stoichiometrically related to the rate of organic carbon

remineralization (e.g., Sayles and Curry, 1988; Ham-

mond et al., 1996). Initially, the consumption rate was

considered to be the sum of two exponential terms,

representing two fractions of remineralized organic

re water d13CT and for d13CPOC

Longitude Latitude Date

57825.25VN 15844.05VW 10/5/98

57824.26VN 15844.32VW 3/7/98

57824.51VN 15844.59VW 2/7/98–3/7/98

57816.92VN 16853.52VW 7/97

57805.30VN 12818.13VW 25/4/98

57805.27VN 12829.45VW 12/5/98

57806.42VN 12829.34VW 29/6/98

56808.52VN 13853.16VW 7/97

52855.11VN 16855.12VW 30/4/98

52854.36VN 16854.05VW 6/7/98

52854.41VN 168 54.57VW 1/5/98–2/5/98

52854.32VN 168 54.35VW 6/7/98–7/7/98

55834.62VN 128 36.06VW 7/97

taken from Black et al. (2001). The sites are arranged by increasing


matter of different reactivity (Berner, 1980; Hammond

et al., 1996). Application of this model to the pore

water O2 profiles yielded unrealistically high depth

attenuation coefficients, such that the most reactive

fraction of organic matter was consumed at x=0. Thus,

a consumption rate with a single exponential term was

adopted, which represents a single fraction of remin-

eralized organic matter (e.g., Hammond et al., 1996).

The steady state mass conservation of pore water O2

in this case is given by Eq. (1) below, where the

quantity in brackets represents measured concentra-

tion (AM), DO2is the bulk sediment diffusion

coefficient (cm2 day�1), x is the depth below the

sediment–water interface (cm), r1 is the O2 con-

sumption rate (AM day�1) at x=0, and l1 is the depth

attenuation coefficient of the consumption rate

(cm�1).

DO2

B2 O2½ �Bx2

¼ r1e�l1x ð1Þ

O2�x ¼ O2�0e�l1x��

ð2Þ

O2½ �0 ¼r1

l21DO2

ð3Þ

IOC ¼ /Z l

0

r1e�l1xdx ¼ /

r1

l1

ð4Þ

The analytical solution of Eq. (1) (Eq. (2)) was

derived by successive integration (Boudreau, 1987)

Table 2

Solute concentration and d13CT value used as upper boundary condition (

Site B Site C

#010/E3 #010/E16

[CT]BWa 2165F14 2183F

d13CT, BWa 0.24F0.08 0.41F

[O2]BW 232.5F7.1 250F[O2]0 227.2 220.7 –

[AT]BWb 2315.5 2307.

XBW, calcite 2.3 1.9

XBW, aragonite 1.6 1.4

The in-situ measured O2 concentration in bottom water ([O2]BW) and at x=

and total alkalinity ([AT]BW) are in Amol kg�1. The carbon isotope ratio of

water saturation with respect to carbonate minerals calcite (XBW, calcite) and

pressure and temperature, and a salinity of 35 p.s.u, using the equations ia Measurements obtained from CTD casts (sites A and C, n=1) and lb Measurements obtained during the METEOR cruise in the proximity

Kiel, Germany, unpublished data).

for known concentration at the sediment–water inter-

face (measured [O2]0, Table 2) as the upper boundary

condition, and zero concentration and concentration

gradient at infinite depth below the sediment–water

interface as the lower boundary condition. The above

formulation is equivalent to that for a first order O2

consumption rate in Rasmussen and Jorgensen (1992).

Curve fitting and parameter prediction were made by

minimizing the residual sum of squares with the

Solver routine on Microsoft Excel, with l1 the fitting

parameter and r1 calculated from Eq. (3). The IOC

can then be computed from Eq. (4). Application of

this model to previously reported O2 micro-electrode

profiles from sediments in the equatorial Atlantic

(stations 13 and 14.1 in Archer et al., 1989) gave

estimates of organic carbon remineralization rates

from Redfield stoichiometric conversion of the

calculated IOC within 100% and 20% of their

reported organic carbon rain rate values.

3.2. Pore water d13CT

The flux of remineralized carbon at the sediment–

water interface was estimated from numerical simu-

lation of the depth distribution of pore water d13CT,

which is based on the steady state mass balance of

Martin et al. (2000) and McArthur (1989). Briefly, the

sediment column is divided in sections as sampled,

and the pore water d13CT within the ith sediment

section (d13CT,i) is considered to result from the

x=0) for the modelling of pore water profiles

Site A

#004/E1 #004/E6 #012/E3

13 2127

0.13 0.20

10 267.5F3.2

254.8 264.0 256.3

3 2344.2

1.5

1.1

0 ([O2]0) is in AM. The bottom water concentration of CT ([CT]BW)

[CT]BW (d13CT, BW) is expressed in x VPDB. The degree of bottom

aragonite (XBW, aragonite) was calculated from CT and AT for in situ

n Keir (1979; 1980) and Millero (1979; 1995) (see text for details).

ander deployments (site B, n=6; site C, n=3).

of sites B and C in May 1997 (Prof. A. Koertzinger, University of


modification of the carbon isotope ratio of bottom

water CT (d13CT,BW) by that of the CT added to the

pore water (d13CT,added) during microbial metabolism

within this section. Thus, the following mass balance

holds:

d13CT;i CT½ �i ¼ �13CT;BW CT½ �BWþ �13CT;added CT½ �i � CT½ �BW

� �: ð5Þ

The d13CT,added depends on the isotopic composi-

tion and relative contribution of the sources of

remineralized carbon to the pore water, i.e., organic

carbon and, if CaCO3 dissolution occurs, carbonate

carbon (Sayles and Curry, 1988). The occurrence of

sedimentary CaCO3 dissolution in deep-sea sediments

depends on the rate and depth scale below the

sediment–water interface of organic matter reminer-

alization by O2, as well as by the ratio of POC to

CaCO3 rain rates (Emerson and Bender, 1981;

Wenzhofer et al., 2001; Pfeifer et al., 2002). It is also

influenced largely by the degree of bottom water

saturation with respect to the carbonate minerals

calcite (XBW,calcite) and aragonite (XBW,aragonite)

(Emerson and Bender, 1981). The degree of saturation

at in-situ temperature and pressure was calculated

using the equation in Keir (1980). The concentration

of carbonate ion in seawater at equilibrium with the

carbonate minerals ([CO32�]sat) was derived using the

equations in Millero (1979, 1995) for an (assumed)

constant Ca2+ concentration at the measured salinity

of 35 p.s.u. (Black et al., 2001). Its concentration in

the bottom water at the BENBO sites ([CO32�]BW)

was calculated from the measured bottom water CT

([CT]BW, Table 2) and total alkalinity (AT) measure-

ments taken at comparable depths in the proximity of

sites B and C during the METEOR cruise (Fig. 1) in

May 1997 (A. Koertzinger, University of Kiel,

Germany, unpublished data), using the equations in

Keir (1979). The calculations indicate that the bottom

water at the BENBO sites is supersaturated with

respect to both carbonate minerals (Table 2) to a

degree and in a decreasing trend with water depth

which agree with that in Wilson and Wallace (1990).

Under these conditions, the excess bottom water

CO32�, D[CO3

2�]BW=[CO32�]BW�[CO3

2�]sat, is an

important agent in the neutralization via acid–base

reactions of the metabolic CO2 produced during

microbial metabolism in the underlying sediments

(Emerson and Bender, 1981; Martin and Sayles,

1996). Based on these calculations, organic matter is

assumed to be the most important source of reminer-

alized carbon to the pore water, and, hence, at steady-

state, the estimated flux of remineralized carbon across

the sediment–water interface is equivalent to the depth-

integrated rate of organic carbon oxidation (Rorg).

For the initial values of [CT]BW and d13CT,BW in

Eq. (5), we use measurements based on CTD casts

and lander deployments (Table 2). The measured

d13CPOC was taken to represent the isotopic compo-

sition of remineralized organic carbon, and, hence,

d13CT,addedid13CPOC. Uncertainty in d13CT,added due

to the reported isotopic fractionation in the order of 1–

3x associated with the oxidation of bulk sedimentary

POC (De Lange, 1998) does not affect the calcu-

lations greatly, generating an uncertainty in the

estimates of Rorg of less than 0.1 mmol m2 day�1.

The average d13CPOC at site C was �20.7F0.1x(n=16), while, at site A, it was �19.1F2.3x (n=13)

in the Holocene layer and �24.3F0.5x (n=13) in the

Glacial layer (i.e., below 13 cm depth). Given the

relatively large variability of the measurements from

the Holocene layer and the possibility of upward

mixing from the Glacial layer below (e.g., Thomson et

al., 2000), both average d13CPOC values are used for

the simulation at site A. This uncertainty in d13CPOC

is translated into ca. 0.1 mmol m�2 day�1 uncertainty

in the estimated Rorg at this site (Table 3). The

d13CPOC at site B was assumed to be similar to that

measured at site C.

The ex-situ measurements of the pore water CT

concentration (unpublished data) were not suitable to

be used in the model because they were affected by

the artefact associated with core recovery (e.g.,

Murray et al., 1980; Martin et al., 2000). The

concomitant decompression effect on pore water

d13CT has been demonstrated to be equivalent to that

induced by calcite precipitation at isotopic equilibrium

with pore water CT and has been found to be

insignificant (i.e., b0.5x) by comparison with the

isotopic change with depth induced by diagenetic

reactions in sediments (McCorkle et al., 1985; Martin

et al., 2000). Because of the decompression artefact,

the profile of pore water CT concentration was

approximated by an empirical exponential function

of depth (e.g., Sayles and Curry, 1988) and was, thus,

set to increase as a result of remineralization reactions

Table 3

Calculations based on model-derived best-fits to in situ pore water

O2 and d13CT profiles

Site Station Profile IOC Rorg l

B 54407#3 d13CT 0.48 0.48

54702#8 d13CT 0.57 0.60

#010/E3 O2 1.48 1.14a 0.84

#010/E16 O2 N0.91 N0.70a 0.10

C 54409#2 d13CT 0.26 0.18

54701#5 d13CT 0.42 0.32

A 54703#1 d13CT 0.36b 0.35

0.29c

#004/E1 O2 0.51 0.37a 0.47

#004/E6 O2 0.48 0.39a 0.47

#012/E3 O2 0.37 0.28a 0.34

The depth-integrated rates of O2 consumption (IOC) and organic

carbon remineralization (Rorg) are in mmol m�2 day�1. The depth

attenuation constant of reaction rates (l) is in cm�1.a Calculated from IOC assuming a Redfield C:O2 ratio of 0.768.b Estimated using d13CPOC=�19x.c Estimated using y13CPOC=�24x.


balanced by diffusion in the sediment from the bottom

water value ([CT]BW) to a steady-state asymptotic

concentration at infinite depth below the sediment–

water interface ([CT]l):

CT½ � ¼ CT½ �l þ CT�BW � CT�l� �

e�l2x��

ð6Þ

where l2=the attenuation coefficient (cm�1) of the

increase in [CT] and the decrease in the remineraliza-

tion rate with depth in the sediment. The average CT

concentration within a sediment section (xi, xj), i.e.,

[CT]i in Eq. (5), is then given by:

CT½ �i ¼

Z xj

xi

CT½ �dx

xj � xi� � ð7Þ

Substitution of Eq. (7) in Eq. (5) and re-arrange-

ment yield d13CT,i as a function of [CT]i, which was

fitted to the pore water d13CT profiles as described

before, with [CT]l and l2 as fitting parameters. The

steady-state CT flux across the sediment–water inter-

face and, hence, Rorg were computed from the CT

gradient at x=0 by taking the first derivative of Eq. (6)

and using Fick’s first law of diffusion (Berner, 1980),

the measured average porosity (see Section 2.1) and

[CT]BW (Table 2), the predicted [CT]l and l2, and the

bulk sediment diffusion coefficient (cm2 day�1) of

dissolved bicarbonate ion (HCO3�), the most abundant

CT species at the pH range of deep-sea sediments.

Application of this model to relevant data in Sayles

and Curry (1988) gave estimates of Rorg within 10–

20% of their published values.

3.3. Diffusion coefficients

All bulk sediment diffusion coefficients were

calculated from the free solution diffusion coefficient

(Do). Specifically, the Do for O2 at 5 8C was taken

from Broecker and Peng (1974), while the Do for all

other solutes from the Handbook of Chemistry and

Physics (Lide, 1994). These values were corrected

for in-situ temperature using the Stokes-Einstein

relationship and Nernst equation (Li and Gregory,

1974), with water density data taken from the

Handbook of Chemistry and Physics (Lide, 1994).

Further correction for sediment tortuosity was made

by using the average / and dividing the temperature-

modified Do by 1�ln(/2) (Boudreau, 1996).

4. Results

4.1. Pore water O2

The in-situ O2 profiles from sites B and A (Figs. 2

and 3) displayed a continuous decrease with depth,

primarily reflecting the utilization of O2 during the

microbial oxidation of organic matter. The sensors

reached the depth of maximum O2 penetration only at

site B at approximately 2.1 and 2.5 cm in the sediment

(Fig. 3; Black et al., 2001). This is considerably

shallower than inferred from the distribution of solid

phase manganese with depth in the sediment at this site

(9 cm; Thomson et al., 2001). Spatial variability, as

well as the difference in the time scale of diagenetic

processes, which are reflected in the instantaneous in-

situ pore water O2 measurements presented here and

the solid phase manganese profiles (several decades;

Thomson et al., 2001), may be responsible for the wide

range of the estimated O2 penetration depth in the

sediments at Site B. The model prediction of the O2

penetration depth at site A, i.e., the depth below which

[O2]b1%[O2]BW (Cai and Sayles, 1996, with [O2]BWgiven in Table 2), was 9.7, 11.1 and 13.4 cm based on

the three available profiles #004/E1, #004/E6 and

#012/E3, respectively. Solid phase geochemical meas-

urements in turbiditic sediment layers within 30–60

Fig. 2. In-situ O2 microelectrode profiles in sediments from site A (Black et al., 2001). Model-derived best-fit curves are based on Eq. (2) (solid

lines).


cm depth below the sediment–water interface at site A

suggest that the extent of the oxic layer can be deeper

at this site (J. Thomson, unpublished data). Based on

the individual IOC estimates from each of the available

pore water O2 profiles (Table 3), the average areal O2

consumption rate in the sediment of site A was

estimated to be 0.45F0.07 mmol m�2 day�1.

There was considerable discrepancy between

observations and the model-generated best fit curves

Fig. 3. In-situ O2 microelectrode profiles in sediments from site B

(Black et al., 2001). Model-derived best-fit curves are based on Eq.

(2) (thin solid line) and on Eq. (8) (thick solid line).

based on Eq. (2) at site B. Specifically, the model

overestimated the gradient in the upper part of both

profiles and, conversely, underestimated the gradient

in their lower part (Fig. 3). The bottom boundary

condition that leads to the fitted Eq. (2) is relaxed,

leading to a slow attainment of zero concentration and

gradient deeper in the sediment than the actual

measurements indicated. Hence, a fixed depth below

the sediment–water interface at which the [O2]

approaches zero appears more appropriate as bottom

boundary condition for the O2 profiles at site B. The

analytical solution of Eq. (1) after modification of its

bottom boundary condition along these lines is given

in Eq. (8) below, with L=depth at which [O2]=0 (taken

from Black et al., 2001), FL=the O2 flux at x=L, l1

and FL being now the fitting parameters, and r1

calculated from the condition of zero concentration at

x=L (Eq. 9).

O2½ �x ¼r1

l21DO2

e�l1x � 1ð Þ

þ r1

l1DO2

e�l1L � FL

/DO2

� �xþ O2�0

�ð8Þ

r1 ¼l21DO2

LFL

/DO2

� O2½ �0� �

e�l1L l1Lþ 1Þ � 1ð ð9Þ

Fig. 4. Profiles of pore water d13CT (xVPDB) from sites B

(squares), C (triangles) and A (circles). The bottom water d13CT

value is indicated by the filled symbols. The horizontal dashed lines

indicate the estimated O2 penetration depth, while the solid lines

indicate model generated curves.


Eq. (8) above approximated the O2 profiles from

site B better than Eq. (2) but failed to reproduce the

irregularities of profile #010/E16 (Fig. 3). While these

can be attributable to excessive electrode noise during

measurement, it is possible that the measurements in

this profile were features of bio-irrigation (e.g., Hales

and Emerson, 1997), which is not accounted for in the

model. Therefore, the IOC estimated from this profile

(Table 3) is a minimum, while the IOC from its

companion and better approximated profile #010/E16

was calculated to be 1.48 mmol m�2 day�1. The flux

of O2 at the base of the modelled layer was predicted

to be N3�10�4 mmol m�2 day�1, i.e., close to zero, in

both cases.

The bottom water O2 concentration at site C (Table

2), measured with a micro-titrator in samples collected

during a lander deployment (Black et al., 2001), is the

only direct measurement available from this site. Both

[O2]BW and sedimentary POC content are major

agents acting on O2 dynamics in the pelagic and

hemipelagic sediments (e.g., Cai and Sayles, 1996;

Cai and Reimers, 1995). The [O2]BW at site C was

higher than that at site B (Table 2), but the similar

sediment organic carbon content at both sites (see

Section 2.1; also, Papadimitriou et al., 2002) suggests

that the O2 penetration depth at site C can be similar to

that at site B (2–2.5 cm; Black et al., 2001). The

thickness of the sedimentary oxic layer at site C was

inferred to be 6.5 cm based on the depth distribution

of solid phase manganese (Thomson et al., 2001).

Pore water NO3� profiles from this site (unpublished

data), which were contemporaneous with the pore

water profiles presented here and similar to those

obtained from site B (shown in Black et al., 2001),

indicated a decline in the concentration of this oxidant

at depths greater than 2–4 cm below the sediment–

water interface, presumably by consumption via post-

oxic organic carbon remineralization (Froelich et al.,

1979). Based on the above, a range of 2–4 cm is

adopted for the O2 penetration depth in the sediment

in subsequent discussion relevant to site C.

4.2. Pore water d13CT

At all three sites, the d13CT exhibited continuous

isotopic depletion with depth in the sediment (Fig. 4),

i.e., it became more negative than the bottom water

value (Table 2) as a result of addition of 13C-depleted


metabolic CO2 to the pore water. At the shallowest

site B, the isotopic depletion relative to bottom water

d13CT was measured to be �1.6x and �1.7x in the

2–3 cm depth layer, where O2 was exhausted (Black

et al., 2001). Thereafter, the profiles showed addi-

tional smooth relative isotopic depletion to a max-

imum of �2.4x at 13–14 cm depth in May and

�2.7x below 10 cm depth in July. At site C, both

profiles exhibited a mean isotopic depletion relative to

bottom water d13CT of �1.7x in the 2–4 cm sediment

layer, in which pore water O2 is inferred to be

depleted (see previous section). Thereafter, the pro-

files exhibited a smooth attainment of a maximum

relative isotopic depletion of �3.7x at 22–23 cm

depth in May and, on average, �2.9x within the 10–

18 cm sediment layer in July. At the deepest site A,

the d13CT profile attained progressively an asymptotic

value which was isotopically depleted by �2.3xrelative to the bottom water value below 10 cm depth,

where O2 became exhausted (see Section 4.1). The

Rorg (Table 3) calculated from the best fit curves on

the measured pore water d13CT profiles (Fig. 4) was in

the narrow range of 0.3–0.6 mmol C m�2 day�1, with

a tendency for higher rates at sites B and C and, also,

slightly higher rates in July at these shallower sites.

Fig. 5. Sediment O2 consumption in eastern north Atlantic

(Porcupine Abyssal Plain, Porcupine Sea Bight, Rockall Trough

and Goban Spur) in different years from 1979 to 1998. Closed

symbols indicate measurements taken in spring (April–May), while

open symbols indicate measurements taken in mid- to late summer

and autumn (July–October): triangles=Lohse et al. (1998), benthic

lander measurements (TOC), and Soetaert et al. (1998), onboard

micro-electrode profiles and diagenetic modelling (IOC); circles=

Pfannkuche (1993), in-situ respirometer measurements (TOC)

squares=Lampitt et al. (1995), in-situ suspended core with onboard

end-point O2 determinations (TOC); filled diamonds=Patching et al

(1986), in-situ suspended core with onboard end-point O2 determi-

nations (TOC); open diamonds=this study, in-situ micro-electrode

profiles and curve fitting (IOC).

5. Discussion

5.1. Sediment O2 consumption in the abyssal and

bathyal eastern north Atlantic

Measurements of sediment O2 consumption in the

eastern north Atlantic have been made with various

techniques since 1979 both within the period of

development and after the collapse of the phytoplank-

ton bloom (Patching et al., 1986; Pfannkuche, 1993;

Lampitt et al., 1995; Lohse et al., 1998). The BENBO

sites were visited within the period of development

and settling of the spring–summer phytoplankton

bloom. Phytodetrital aggregates were recorded on

the sea floor in July 1998 at the shallowest site B.

Aggregate deposits of varying thickness (up to 2 cm)

were recovered in megacores from site C during both

cruises in 1998, while, at the deepest site A,

phytodetrital aggregates were not observed on any

of the cruises. The IOC estimates in the sediments

(Table 3) fall within the lower end of those reported

for spring and early summer in the general area and

are very much lower than the maxima recorded in late

summer and early autumn in response to the settling

of the spring–summer phytoplankton bloom (Fig. 5).

The O2 data set used here is not extensive and, as

such, can not resolve the issue of seasonality in

benthic metabolism in the year 1998. Comparison

with the temporally extended measurements of the

past studies in the area suggests that the activity of the

microbial sediment community had not responded by

the beginning of July 1998, even though phytodetritus

had been observed and recovered in multicores at sites

B and C. Phytodetrital aggregates can be heavily

utilized above the sediment by microbial oxidation

(Lochte and Turley, 1988; Smith et al., 1996) and

consumption by epibenthic fauna (Beaulieu, 2002,

and references therein). Furthermore, a measurable

response to this seasonal pulse of organic matter will

be manifest in pore waters only if this material is

rapidly mixed into the sediment surface and is highly

;

.


reactive upon incorporation (Martin and Bender,

1988; Sayles et al., 1994). It is possible, therefore,

that the phytodetrital aggregates sighted during

sampling were transient or consisted primarily of

heavily degraded material which would have little

impact on the sediment microbial activity and, hence,

IOC upon slow mixing into the underlying sediments.

The above comparisons are tentative because the

majority of the previously reported measurements in

the area were obtained by benthic landers and

respirometers and, hence, represent total sediment

community O2 consumption rates (TOC), which are

often higher than IOC estimates obtained from pore

water profiles (Wenzhofer and Glud, 2002). The

discrepancy noted here, therefore, may be attributable

to this phenomenon.

5.2. Organic carbon remineralization

The maximum isotopic depletion measured in the

sediment from the deepest site was attained in the

lower part of the oxic zone (Fig. 4), indicating the

dominant role of O2 consumption in organic carbon

remineralization at this site. This is supported by the

good agreement of the Rorg predicted from the d13CT

profile with that estimated from the in-situ O2 profiles

from this site (Table 3). In contrast, 30–50% of the

maximum isotopic depletion at the two shallower sites

was measured well below the oxic layer (i.e., b2 cm

depth; Fig. 4). This suggests occurrence of post-oxic

organic carbon remineralization in the sediments of

these sites. Occurrence of post-oxic metabolic pro-

cesses is supported by the exhaustion of pore water

NO3� by 10 cm depth below the sediment–water

interface (e.g., Black et al., 2001) and the develop-

ment of deeper (N10 cm depth) DOC concentration

maxima at sites B and C (Papadimitriou et al., 2002),

which are typical of anoxic metabolism (e.g., Burdige,

2002). The profiles of bulk sedimentary organic

carbon from these sites showed deep concentration

maxima below 6–10 cm depth, after an apparent

exponential decrease (Thomson et al., 2000; Papadi-

mitriou et al., 2002), with 45% to ~100% of the

decrease occurring the upper 3–5 cm of the sediment,

i.e., where O2 was estimated to be utilized to

exhaustion. The deep POC concentration maxima

were attributed to non-local particle transport from the

sediment surface by infaunal organisms (Thomson et

al., 2000). Biotic non-local POC transport from the

surface to sediment layers below the oxic layer can be

conceived to facilitate anoxic processes in these deep-

sea locations by providing suitable organic substrate

to anaerobes, much fresher and, hence, more labile

than the relict organic matter that would reach these

sediment layers at steady state after extensive oxic

remineralization via the slow sedimentation and

biodiffusive sediment mixing measured at these sites.

Despite indications for contribution of post-oxic

organic carbon remineralization processes, the Rorg

calculated for site B from the d13CT profiles is lower

by a factor of 2–3 than that estimated from the in-situ

O2 profiles, which indicate an organic carbon remi-

neralization rate in the order of 1 mmol m�2 day�1 or

higher (Table 3). Oxygen consumption by the

oxidation of reduced products of post-oxic metabo-

lism diffusing from deeper sediment layers rather than

directly by organic carbon oxidation cannot explain

the discrepancy because the former process is

indirectly linked stoichiometrically to the latter. The

estimates of Rorg from the d13CT profiles were based

on the assumption that organic carbon is the sole

source of CT added to the pore water during sediment

metabolism. Dissolution of sedimentary CaCO3 can

result from the acidification of pore water during

diagenetic redox reactions in the presence of O2 and is

thus related to the rate of O2 consumption near the

sediment surface (Emerson and Bender, 1981; Bou-

dreau, 1987; Archer et al., 1989; Hales and Emerson,

1997; Martin et al., 2000; Wenzhofer et al., 2001;

Pfeifer et al., 2002). While the post-depositional cycle

of CaCO3 minerals may involve initially a precip-

itation stage in the surface of CaCO3-rich sediments

underlying supersaturated oceanic waters (Jahnke and

Jahnke, 2004), recent in-situ investigations in deep-

sea locations in the Atlantic (N1000 m water depth)

demonstrated occurrence of CaCO3 dissolution in

sediments overlain by moderately to substantially

supersaturated bottom waters (Martin and Sayles,

1996; Wenzhofer et al., 2001; Pfeifer et al., 2002). It

is possible, therefore, that despite conditions of

carbonate mineral supersaturation of the bottom water

at the BENBO sites (Table 2), the available excess

carbonate ion may be insufficient for the neutraliza-

tion of the metabolic CO2 produced in the oxic layer

of the underlying sediments and some CaCO3

dissolution may still occur. The stable carbon isotope

Fig. 6. d13CT (in xVPDB) measured in the base of the oxic

sediment layer (open symbols) and in the deepest sediment layers

sampled below the oxic layer (filled symbols) vs. DO2i[O2]BW in

sites B (squares), C (triangles) and A (circles). Stars indicate the

d13CT values at the bottom of the oxic layer predicted for the

measured d13CT, [CT] and [O2] in the bottom water, using numerica

stoichiometric modelling for O2 consumption only with (dashed

lines) and without CaCO3 dissolution (solid lines). The dashed and

solid lines do not represent functional trends, while vertical error

bars indicate the range of measurements within the sediment layer

where pore water [O2] ~0.


ratio of bulk CaCO3 (d13CCaCO3) in the upper 40 cm

of the sediments averaged +0.25F0.36x,+0.58F0.15x and+0.49F0.37x at sites B, C and A,

respectively (Thomson et al., 2000). If CaCO3

dissolution occurred, the resulting carbonate carbon

added to the pore water by this process would be

isotopically more positive than that added by the

oxidation of organic carbon by ca. 20x. Incorpora-

tion of an isotopically enriched CT source into the

carbon mass balance would shift the d13CT predicted

in each sediment section towards isotopically enriched

values, depending on the rate of dissolution relative to

that of organic carbon oxidation by O2 consumption

(e.g., McArthur, 1989). In this case, a higher Rorg

would be required to fit the measurements than that

predicted assuming no CaCO3 dissolution, all other

parameters considered adequately constrained. Hence,

the current Rorg estimates from the pore water profiles

of d13CT at these sites assuming organic carbon

remineralization only (Table 3) are likely to be a

minimum.

The potential for CaCO3 dissolution within the

oxic layer and the intensity of post-oxic metabolic

processes can be explored by comparing the d13CT

measurements at the base of the oxic layer and in the

deepest sediment sections sampled with the d13CT that

can be predicted as a function of bottom water O2

concentration, using a stoichiometric numerical model

for oxic remineralization only with and without

occurrence of CaCO3 dissolution (McCorkle and

Emerson, 1988; McArthur, 1989). In this case, the

pore water d13CT at depth in the sediment where

[O2]~0 can be computed from Eq. (5) by predicting

DCTi[CT]i�[CT]BW from the following equation

(McCorkle and Emerson, 1988):

DCT ¼ b 1þ �ð Þ DO2

DHCO�3

DO2 ð10Þ

where DO2i[O2]BW (Table 2), b is the molar ratio of

organic carbon remineralized to O2 consumed during

oxic metabolism, a is the stoichiometric ratio of the

rates of CaCO3 dissolution to oxic organic carbon

remineralization, and DO2, DHCO3

� are the bulk sedi-

ment diffusion coefficients of O2 and the bicarbonate

ion. Subsequent application of the isotopic mass

balance (Eq. (5)) requires the calculation of d13CT,added

for two sources of pore water CT, i.e., organic carbon,

with a d13CPOC as before (see Section 3.2), and

carbonate carbon with a d13C equivalent to that of

bulk sedimentary CaCO3 (d13CCaCO3) given above.

The d13CT,added is then given by the equation

(McArthur, 1989):

d13CT;added ¼ad13CCaCO3

þ d13CPOC

1þ að11Þ

The predicted pore water d13CT using the above

rationale was computed assuming that the reactions

proceed with Redfield stoichiometry, hence, b=0.768and a=1.06 or a=0 for oxic remineralization with and

without CaCO3 dissolution respectively (e.g., McAr-

thur, 1989). Comparison of the actual d13CT measure-

ments with the results of these calculations helps

illustrate the following points.

The mean of the observations from the base of the

oxic layer at sites B and C fall within V0.1x of the

predicted values for oxic remineralization with CaCO3

dissolution (Fig. 6). The range of these measurements,

however, was higher than the resolution offered by the

model for the occurrence or not of dissolution,

especially so at site C, where the uncertainty of the

O2 penetration depth is also large. The isotopic

l


depletion measured in deeper sediment layers well

below the oxic layer at sites B and C (Fig. 6) was

substantial relative to the measurements at the base of

the oxic layer and the predictions for oxic remineral-

ization alone (ca. 1–2x difference). This demon-

strates occurrence of post-oxic organic carbon

remineralization, while the magnitude of the isotopic

depletion of pore water CT in deep sediment sections

suggests that post-oxic remineralization should be

more intense at site C than at site B (Fig. 6). In

contrast, the mean of the observations from the base

of the oxic layer at site A (i.e., between 10 and 13 cm

depth) can be closely predicted for organic carbon

remineralization only with a d13CPOCi�24x, while

the subsequent isotopic depletion of the pore water CT

was comparatively negligible (approximately �0.2x;

Fig. 6). This indicates dominance of oxic remineral-

ization and minimal contribution of post-oxic pro-

cesses at the deepest site. It further illustrates the

internal consistency of the dataset from this site, with

the matching Rorg estimates from the pore water d13CT

and O2 (Table 3) within a ~30% uncertainty in Rorg

associated with a ~ 5x uncertainty in the actual d13Cof the remineralized organic carbon.

Post-oxic remineralization adds metabolic CO2 with

a d13Cid13CPOC (e.g., McCorkle and Emerson, 1988)

and causes the pore water d13CT measurements at the

base of the oxic layer to be more depleted isotopically

than in the case of oxic remineralization alone. This

offset was shown to be a function of O2 depth

penetration below the sediment–water interface and

intensity of post-oxic remineralization (McCorkle and

Emerson, 1988). For the range of estimates of the O2

penetration depth at sites B and C (i.e., between 2 and 4

cm depth), this offset was found to be up to 0.5x(op. cit.). Whereas these considerations put the observ-

ations from the base of the oxic layer at sites B and C

more firmly in the region of occurrence of CaCO3

dissolution (Fig. 6), this process remains only a possi-

bility and cannot be ascertainedwith the present data set.

6. Conclusions

Pore water profiles of O2 (in-situ measurements)

and of the stable isotopic composition of total

dissolved inorganic carbon (ex-situ measurements)

provided estimates of the depth-integrated rates of

diffusion-supplied O2 consumption (IOC) and organic

carbon remineralization (Rorg) respectively during

microbial metabolism in sediments at 1100, 2000 and

3500 m water depth in the eastern north Atlantic.

Stoichiometric and isotopic constraints indicated that

oxic oxidation of organic carbon was the most

important pathway of sedimentary carbon reminerali-

zation at 3500 m water depth. At this site, the ratio of

the independently estimated Rorg (ca. 0.3 mmol C m�2

day�1) to the average IOC (0.45F0.07 mmol O2 m�2

day�1) was 0.6–0.8, which is consistent with the

Redfield stoichiometry for the oxidation of organic

matter (i.e., 0.768). In contrast, stoichiometric and

isotopic constraints indicated occurrence of post-oxic

remineralization of organic carbon in the sediments

from the two shallower sites. This is coincident with

biological sediment structures and sub-surface (N10

cm) maxima in the profiles of dissolved and particulate

organic carbon and suggests a link of post-oxic

metabolism to infaunal activity. The Rorg (0.5–0.6

mmol C m�2 day�1) to IOC (1–1.5 mmol O2 m�2

day�1) ratio at the shallowest site (~0.3–0.4) was lower

than Redfield stoichiometry. The occurrence of meta-

bolically induced CaCO3 in the sediments, though not

improbable, cannot be documented unambiguously

from the present data set and was not taken into account

in estimating Rorg. This uncertainty is relevant to the

isotopic mass balance of remineralized carbon, and the

current Rorg derived from pore water d13CT profiles

may thus be an underestimate of their true value.

Acknowledgements

We thank the masters and crew members of RRS

Charles Darwin for their help during cruises. We

thank G. Fones and K. Black for providing the oxygen

microelectrode data, D. Hughes, Dunstaffnage Marine

Laboratory, UK, for providing information on animal

abundances and burrows, and P. Kennedy for his

excellent technical assistance. We are grateful to Prof.

A. Koertzinger, University of Kiel, Germany, for

allowing us to access unpublished data from the

METEOR cruises in the eastern north Atlantic.

Finally, we thank Drs. G.J. De Lange, John Thomson

and D. Archer for their constructive review of the

manuscript. This work was funded by NERC grants

GST/02/1754 and GR3/EOO68.


References

Archer, D., Emerson, S., Reimers, C., 1989. Dissolution of calcite in

deep-sea sediments: pH and O2 microelectrode results. Geo-

chimica et Cosmochimica Acta 53, 2831–2845.

Barnett, P.R.O., Watson, J., Connelly, D., 1984. A multiple corer for

taking virtually undisturbed samples from shelf, bathyal and

abyssal sediments. Oceanologica Acta 7, 399–408.

Beaulieu, S.E., 2002. Accumulation and fate of phytodetritus on the

sea floor. Oceanography and Marine Biology: an Annual

Review 40, 171–232.

Berner, R.A., 1980. Early Diagenesis: A Theoretical Approach.

Princeton University Press, Princeton, NJ.

Black, K.S., Fones, G.R., Peppe, O.C., Kennedy, H., Bentaleb, I.,

2001. An autonomous benthic lander: preliminary observations

from the UK BENBO thematic programme. Continental Shelf

Research 21, 859–877.

Boudreau, B.P., 1987. A steady-state diagenetic model for dissolved

carbonate species and pH in the pore waters of oxic and suboxic

sediments. Geochimica et Cosmochimica Acta 51, 1985–1996.

Boudreau, B.P., 1996. The diffusive tortuosity of fine-grained

unlithified sediments. Geochimica et Cosmochimica Acta 60,

3139–3142.

Broecker, W.S., Peng, T.H., 1974. Gas exchange rates between air

and sea. Tellus 26, 21–35.

Burdige, D.J., 2002. Dissolved organic matter in sediment pore

waters. In: Hansell, D., Carlson, C. (Eds.), Biogeochemistry of

Marine Dissolved Organic Matter. Academic Press, New York,

pp. 611–663.

Cai, W.J., Reimers, C.E., 1995. Benthic oxygen flux, bottom water

oxygen concentration and core top organic carbon content in the

deep northeast Pacific Ocean. Deep-Sea Research. Part 1.

Oceanographic Research Papers 42, 1681–1699.

Cai, W.J., Sayles, F.L., 1996. Oxygen penetration depths and fluxes

in marine sediments. Marine Chemistry 52, 123–131.

De Lange, G.J., 1998. Oxic vs. anoxic diagenetic alteration of

turbiditic sediments in the Madeira Abyssal Plain, eastern north

Atlantic. In: Weaver, P.P.E., Schmincke, H.-U., Duffield, E.

(Eds.), Proceedings of the Ocean Drilling Program, Scientific

Results 157, 573–579.

Emerson, S.R., Bender, M., 1981. Carbon fluxes at the sediment–

water interface of the deep-sea: calcium carbonate preservation.

Journal of Marine Research 39, 139–162.

Froelich, P.N., Klinkhammer, G.P., Bender, M.L., Luedtke, N.A.,

Heath, G.R., Cullen, D., Dauphin, P., Hammond, D., Hartman,

B., Maynard, V., 1979. Early oxidation of organic matter in

pelagic sediments of the eastern equatorial Atlantic: Suboxic

diagenesis. Geochimica et Cosmochimica Acta 43, 1075–1090.

Hales, B., Emerson, S., 1997. Calcite dissolution in sediments of the

Ceara Rise: In situ measurements of pore water O2, pH, and

CO2 (aq). Geochimica et Cosmochimica Acta 61, 501–514.

Hammond, D.E., McManus, J., Berelson, W.M., Kilgore, T.E.,

Pope, R.H., 1996. Early diagenesis of organic material in

equatorial Pacific sediments: Stoichiometry and kinetics. Deep-

Sea Research. Part 2. Topical Studies in Oceanography 43,

1365–1412.

Hofmann, M., Wolf-Gladrow, D.A., Takahashi, T., Sutherland, S.C.,

Six, K.D., Maier-Reimer, E., 2000. Stable carbon isotope

distribution of particulate organic matter in the ocean: a model

study. Marine Chemistry 72, 131–150.

Jahnke, R.A., Jahnke, D.B., 2004. Calcium carbonate dissolution in

deep sea sediments: reconciling microelectrode, pore water and

benthic flux chamber results. Geochimica et Cosmochimica

Acta 68, 47–59.

Keir, R.S., 1979. The calculation of carbonate ion concentration

from total CO2 and titration alkalinity. Marine Chemistry 8,

95–97.

Keir, R.S., 1980. The dissolution kinetics of biogenic calcium

carbonate in seawater. Geochimica et Cosmochimica Acta 44,

241–252.

Kennedy, H., Gacia, E., Kennedy, D.P., Papadimitriou, S., Duarte,

C.M., 2004. Organic carbon sources to SE Asian coastal

sediments. Estuarine, Coastal and Shelf Science 60, 59–68.

Lampitt, R.S., Raine, R.C.T., Billett, D.S.M., Rice, A.L., 1995.

Material supply to the European continental slope: a budget

based on benthic oxygen demand and organic supply. Deep-

Sea Research. Part 1. Oceanographic Research Papers 42,

1865–1880.

Li, Y.H., Gregory, S., 1974. Diffusion of ions in sea water and in

deep sea sediments. Geochimica et Cosmochimica Acta 38,

703–714.

Lide, D.R. (Ed.), Handbook of Chemistry and Physics. CRC Press,

London, pp. 5-90–5-92.

Lochte, K., Turley, C.M., 1988. Bacteria and cyanobacteria

associated with phytodetritus in the deep sea. Nature 333,

67–69.

Lohse, L., Helder, W., Epping, E.H.G., Balzer, W., 1998. Recycling

of organic matter along a self-slope transect across the N.W.

European Continental Margin (Goban Spur). Progress in

Oceanography 42, 77–110.

Martin, W.R., Bender, M.L., 1988. The variability of benthic fluxes

and sedimentary remineralization rates in response to seasonally

variable organic carbon rain rates in the deep sea: A modelling

study. American Journal of Science 288, 561–574.

Martin, W.R., Sayles, F.L., 1996. CaCO3 dissolution in sediments of

the Ceara Rise, western equitorial Atlantic. Geochim. Cosmo-

chim. Acta 60, 243–263.

Martin, W.R., McNichol, A.P., McCorkle, D.C., 2000. The

radiocarbon age of calcite dissolving at the sea floor: estimates

from pore water data. Geochimica et Cosmochimica Acta 64,

1391–1404.

McArthur, J.M., 1989. Carbon isotopes in pore water, calcite and

organic carbon from distal turbidites of the Madeira Abyssal

Plain. Geochimica et Cosmochimica Acta 53, 2997–3004.

McCorkle, D.C., Emerson, S., 1988. The relationship between

pore water carbon isotopic composition and bottom water

oxygen concentration. Geochimica et Cosmochimica Acta 52,

1169–1178.

McCorkle, D.C., Emerson, S.R., Quay, P.D., 1985. Stable Carbon

isotopes inmarine porewaters. Earth Planet. Sci. Lett. 74, 13–26.

Millero, F.J., 1979. The thermodynamics of the carbonate system in

seawater. Geochimica et Cosmochimica Acta 43, 1651–1661.


Millero, F.J., 1995. Thermodynamics of the carbon dioxide system

in the oceans. Geochimica et Cosmochimica Acta 59, 661–677.

Murray, J.W., Emerson, S., Jahnke, R., 1980. Carbonate saturation

and the effect of pressure on the alkalinity of interstitial waters

from the Guatemala Basin. Geochimica et Cosmochimica Acta

44, 963–972.

Papadimitriou, S., Kennedy, H., Bentaleb, I., Thomas, D.N., 2002.

Dissolved organic carbon in sediments from the eastern north

Atlantic. Marine Chemistry 79, 37–47.

Papadimitriou, S., Kennedy, H., Kattner, G., Dieckmann, G.S.,

Thomas, D.N., 2004. Experimental evidence for carbonate

precipitation and CO2 degassing during sea ice formation.

Geochimica et Cosmochimica Acta 68, 1749–1761.

Patching, J.W., Raine, R.C.T., Barnett, P.R.O., Watson, J., 1986.

Abyssal benthic oxygen consumption in the northeastern

Atlantic: measurements using the suspended core technique.

Oceanologica Acta 9, 1–17.

Pfannkuche, O., 1993. Benthic response to the sedimentation of

particulate organic matter at the BIOTRANS station, 478N,208W. Deep-Sea Research. Part 2. Topical Studies in Ocean-

ography 40, 135–149.

Pfeifer, K., Hensen, C., Adler, M., Wenzhofer, F., Weber, B.,

Schulz, H.D., 2002. Modelling of subsurface calcite

dissolution, including the respiration and reoxidation pro-

cesses of marine sediments in the region of equatorial

upwelling off Gabon. Geochimica et Cosmochimica Acta 66,

4247–4259.

Rasmussen, H., Jorgensen, B.B., 1992. Microelectrode studies of

seasonal oxygen uptake in a coastal sediment: role of molecular

diffusion. Marine Ecology. Progress Series 81, 289–303.

Sayles, F.L., Curry, W.B., 1988. d13C, TCO2, and the metabolism of

organic carbon in deep sea sediments. Geochimica et Cosmo-

chimica Acta 52, 2963–2978.

Sayles, F.L., Martin, W.R., Deuser, W.G., 1994. Response of

benthic oxygen demand to particulate organic carbon in the deep

sea near Bermuda. Nature 371, 686–689.

Smith, C.R., Hoover, D.J., Doan, S.E., Pope, R.H., DeMaster, D.J.,

Dobbs, F.C., Altabet, M.A., 1996. Phytodetritus at the abyssal

seafloor across 108 of latitude in the central equatorial Pacific.

Deep-Sea Research. Part 2. Topical Studies in Oceanography

45, 1309–1338.

Smith Jr., K.L., Baldwin, R.J., Glatts, R.C., Kaufmann, R.S., Fisher,

E.C., 1998. Detrital aggregates on the sea floor: Chemical

composition and aerobic decomposition rates at a time-series

station in the abyssal NE Pacific. Deep-Sea Research. Part 2.

Topical Studies in Oceanography 45, 843–880.

Soetaert, K., Herman, P.M.J., Middelburg, J.J., Heip, C., 1998.

Assessing organic matter mineralization, degradability and

mixing rate in an ocean margin sediment (Northeast Atlantic) by

diageneticmodelling. Journal of Marine Research 56, 519–534.

Thomson, J., Brown, L., Nixon, S., Cook, G.T., MacKenzie, A.B.,

2000. Bioturbation and Holocene sediment accumulation fluxes

in the north–east Atlantic Ocean (Benthic Boundary Layer

experiment sites). Marine Geology 169, 21–39.

Thomson, J., Nixon, S., Croudace, I.W., Pedersen, T.F., Brown, L.,

Cook, G.T., MacKenzie, A.B., 2001. Redox-sensitive element

uptake in north-east Atlantic Ocean sediments (Benthic Boun-

dary Layer Experiment sites). Earth and Planetary Science

Letters 184, 535–547.

Wenzhofer, F., Glud, R.N., 2002. Benthic carbon mineralization in

the Atlantic: a synthesis based on in situ data from the last

decade. Deep-Sea Research. Part 1. Oceanographic Research

Papers 49, 1255–1279.

Wenzhofer, F., Adler, M., Kohls, O., Hensen, C., Strotmann, B.,

Boehme, S., Schulz, H.D., 2001. Calcite dissolution driven by

benthic mineralization in the deep-sea: In situ measurements of

Ca2+, pH, pCO2 and O2. Geochimica et Cosmochimica Acta 65,

2677–2690.

Wilson, T.R.S., Wallace, H.E., 1990. The rate of dissolution of

calcium carbonate from the surface of deep-ocean turbidite

sediments. Philosophical Transactions of the Royal Society of

London. A 331, 41–49.

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