Organic matter diagenesis in the northeast Pacific: transition from aerobic red clay to suboxic hemipelagic sediments

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Organic matter diagenesis in the northeast Pacific: transition from aerobic red clay to suboxic hemipelagic sediments .IAM~W. MUngA'C"and KA'n-mwM. Ktnvu~*'l" (Received 21 December 1988;/n revisedform 27 ./u/y 1989; accepted 3 August 1989) Abstract--Analyses for dissolved oxygen, nitrate, silicate and manganese in the interstitial water born an/n s/tu sampler and from boxcore sediment samples have been combined with sofid phase sediment analyses of carbon and nitrogen to study the transition from aerobic to suboxic diagenesl¢ in the northeast Pacific. The station locations coincide with the VERTEX sediment trap stations. This has enabled us to study diagenesis as a function of the flux of organic carbon to the sediment-water interface. Organic carbon in the sediments decreases with distance from the continental margin. At all stations, except 183-2 on the Monterey deep sea fan, organic carbon decreases rapidly below the sediment-water interface. The organic matter at Stas 183-3, 4, 5 and 6 has a C/N molar ratio of 10. At Sta. 183-2 the organic matter has a relatively nitrogen-poor C/N ratio of 15.8, suggesting terrestr/al input. The stoichiometry of the decomposing organic matter at Stas 183-3, 4, 5 and 6 was determined from the porewater oxygen and nitrate, and the resulting C/N ratio was 14. The flux of oxygen into the sediments decreases with increasing water depth. Comparison of the oxygen consumption by the sediments with the rain rate of organic carbon indicates that most of the rain of organic carbon is oxidized within the sediments using oxygen as the electron acceptor. The first order degradation rate constant for organic carbon oxidation calculated from the porewater oxygen profiles was found to correlate strongly with the rain rate of organic carbon. The rate constant increases with the rain rate, suggesting that at higher rain rates the organic matter is more "reactive". The mean life for sedimentary organic carbon is 12,000 years in the central North Pacific and decreases to 200 years near the continental boundary.

INTRODUCTION

AEROBICrespiration is the principal characteristic of the early diagenesis in sediments of the central Pacific (MURRAYand GRUNDMANIS,1980; GRUNDMANISand MURRAY, 1982; JAHNKEet al., 1982a; REIMERSet al., 1984; REIMERS,1987). In sediments from the central and eastern equatorial Pacific (EMERSONet al., 1980; JAHNKEet al., 1982b; BENDERand HEOOIE, 1984; REIMERSand SMrm, 1986) dissolved oxygen is rapidly consumed and the characteristic features of suboxic diagenesis (e.g. denitrification and Mn, Fe reduction) become predominant. Although oxygen is the main electron acceptor for organic carbon oxidation, the secondary oxidants may be important in regulating the burial of organic carbon under conditions of high organic carbon rain rates (BENDERand HEOOIE, 1984). Other than the porewater nitrate, manganese and iron data from the Baja transect at 25~I reported by J ~ et al. (1982), SAWLANand MURRAY(1983) and MURRAYet al. (1984), little is known about the transition from oxic to suboxic diagenesis in the Pacific. * School of Oceanogrsphy, University of Washington, Seattle, WA 98195, U.S.A. t Present address: U.S. Geological Survey, Water Resources Division, Room W-2234 Federal Building, Sacramento, CA 95825, U.S.A. 59

60

J . W . MURRAYand K. M. Kuivtta

JAHNKE and JACKSON(1987a) have shown that benthic respiration along the continental margins in the 3000--4000 m depth range is a major contribution to deep sea metabolism. These calculations depend on the variation of oxygen consumption with depth. Unfortunately, data are largely absent from this critical depth range in the North Pacific. In order to understand the regional and paleo-variation in the organic carbon content of marine sediments it is important to understand the variables that control the rate of organic matter diagenesis and organiccarbon burial. Models that couple organic carbon and oxygen reactions with the rain rate of organic carbon suggest that the master variables for explaining the organic carbon content and diagenesis of sediments are the flux of particulate organic carbon (POC) to the sediments, bottom water oxygen content and the organic matter degradation rate constant (BENDER and HEGOIE, 1984; EMERSON, 1985; EMERSONet al,, 1985). The first order rate constant for organic carbon oxidation can increase with the rain rate of POC, although the origin of the relationship is unclear. In order to address the regional (and water column depth) variation and the relationship of the rate of organic carbon decomposition to the rain rate of POC we collected interstitial water and sediment samples from under the transect of VERTEX sediment trap stations off the west coast of North America (-35°N). The rain rate of POC varies significantly along this transect (MARTIN et aL, 1987) while bottom water oxygen is essentially constant (BRoECKER et al., 1982). This enabled us to eliminate oxygen as a variable and to focus on the influence of the organic carbon rain rate on the diagenesis of organic matter. SAMPLING

AND

ANALYTICAL

METIIODS

All samples were collected during R.V.T.G. Thontpson cruise 183 in September 1984. Our sampling locations were chosen to coincide with the stations of the VERTEX Monterey to tlawaii transect (Fig. I, Table 1). Sta. 183-2 was located at 357{) m on the foot of the Monterey dccp sea fan. Stations 183-3 to 183-6 progressed from hemipelagic

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The station coordinates and water depths are given in Table 1. Also shown are other stations

referred to in the text. These include GEOSECS 201,202 and 204 and site MPG-I.

Organic matter diagenesis

61

Table 1. Station locations and water depth of sediment and interstitial water samples collected during TT183. Also shown are the corresponding VERTEX stations and closest GEOSECS stations

Sta. no. T1"183-2 TT183-3 T1"183-4 TT183-5 TT183-6

Lat. 35"39.9'N 35"24.2'N 35"00.7'N 33"01.9'N 28*01.9'N

Long. 122"32.3'W 124"30.0'W 127"57.2'W 139"30.8'W 155"01.9'W

Water d e p t h (m) 3570 4211 4772 4980 5668

aosest VERTEX Sta. no. V-I V-2 V--4 IV-5

Closest GEOSECS Sta. no. 201 201 202. 204

to red clay sediments, all with essentially no CaCO3, and covered the depth range of 4211-5668 m. The sediments in this central gyre region are oxidized, eolian-rich (quartz and illite) clays (REx and GOLDBERG, 1958). Data from GEOSECS Stas 201,202 and 204, and the mid-Pacific gyre site, MPG-I, are used in the text and their locations are shown in Fig. 1. Interstitial water samples were collected using an in situ sampler and by centrifuging boxcore sediment samples at in situ temperature. The in situ interstitial water sampler was designed to obtain uncontaminated samples for gas analyses and was used as described by GRUNDMANISand MURRAY(1982). A short probe was used that had sampling ports at nominal depths from 2 to 54 cm (at 3 or 5 cm intervals). This sampler provides an in situ sample which is essential for constituents like alkalinity (MURRAYel al., 1980). In return it sacrifices depth resolution, which is especially important near the sedimentwater interface. A large wooden wheel "Frisbee" was used to help support the sampler at the sediment-water interface. We used a Soutar boxcore to collect undisturbed samples near the sediment-water interface (JAIlNKEet al., 1986). All samples to be discussed in this paper were judged at the time of recovery to have a well-preserved sediment-water interface. The overlying water was clear and intact biological structures were normally observed at the sediment surface. The core lengths ranged up to 40 cm. The sediments were gently sub-sampled immediately after retrieval with P V C core liner and transferred to a refrigerated lab van. The samples were allowed to equilibrate at in situ temperature. The cores were then extruded, packed in centrifuge tubes and centrifuged at 16,000 rpm for 5-10 min to extract the interstitial water. The sediments were sampled at 0.5 cm intervals from 0 to 5 cm and 1.0 cm intervals from 5 to 10 cm. Two cm intervals were used at depths >10 cm. The in situ interstitial water samples were analysed soon after recovery for dissolved oxygen, nitrate, silica and manganese. The Soutar core samples were analysed for nitrate, silica and manganese. The redundancy in the analyses provided a check of the profiles in the region where they overlapped. Dissolved Oz + Ar was analysed using 5 ml samples as described in GRUNDMANISand MURRAY(1982). A correction for Ar was made based on the solubility of Ar in seawater at the in situ temperature. Nitrate and silica were analysed manually using micro-methods modified from STmCKLANDand PARSONS (1972). Manganese was analysed by the formaldoxime colorimetric method of BREWER and SPENCER(1974). The analyses of the Soutar core sediment samples included porosity (sea-salt corrected) and total and organic carbon and nitrogen by the liquid acidification methods of GRUNDMANISand MURRAY(1982) and HEDGES and STERN (1984).

62

J . w . Mummy and K. M. Kurvu.A

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Fig. 3. Total carbon, organic carbon and inorganic carbon (CaC03) from Souta,r sediment

cores collected during TYI83. (a) SUB 2-I and 2-2, (b) Stas 3-I, (¢) Stas 4-I and 4-2, (d) Stas 5-I and 5-2, and (c) Slas 6-I and 6-2.

0.5 I

J . W . Mum~v and K. M. KUWU.A

64

RESULTS

The results of the porosity, organic carbon, CaCO3 and total nitrogen analyses are shown in Figs 2 and 3. In all cores a porosity gradient was observed in the top 5 cm (10 era for Sta. 2). The porosity typically decreases by about 5%. On a regional basis, the porosity decreases from a range of 77-83% in core 2 to a range of 66-74% in core 6. The variable porosity in core 2 is probably due to the fact that this core was on the Monterey deep sea fan where sediment accumulation is highly variable. The organic carbon content progressively decreases from about 3% at Sta. 2 to 0.20.3% at Sta. 6 (Fig. 3). At Stas 3-6 organic carbon is sharply higher in the surface samples (usually 0--0.5 cm). The replicate analyses tend to agree within +4% of the measured value. Total nitrogen values are extremely low. At Sta. 2 they decrease from 0.323 to 0.108%. The range at Sta. 6 is 0.072-0.055%. We determined both total carbon and organic carbon on many of the samples. Here we define inorganic carbon as the difference between total carbon and organic carbon. This inorganic carbon is probably a carbonate phase. The %C as inorganic carbon is about 0.1% or less at all stations except Sta. 2, where it averages about 0.4%. T T183-STN 2

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Fig. 4. Interstitial water profiles of oxygen, nitrate, manganese and silica at Sta. 2 of TT183. The following data are shown: oxygen, below detection in all samples; nitrate, one boxcore; manganese, one boxcore and one in sit-,; silica, one boxcore and two/n $itu. Arrows indicate the bottom water concentrations either estimated from our bottom water samples from the/n $itu sampler or from the closest GEOSECS station. All concentrations are in pmoi kg -t. Results are shown for both the in sit-, sampler and from Soutar boxcores. The/n s/tu samples are indicated by the closed circles and x's that correspond to the distance from the base plate to the middle of the sample port on the probe. Observation of the mud packs that are often recovered on the probe of the/n sit,, sampler suggests that the samples are probably drawn from a hemispherical region with a radius of 2-3 era. The Soutar core samples are represented by vertical bars or open circles corresponding to the thickne~ of the interval sampled.

Organic matter diagenesis

65

At Sta. 183-2, oxygen was below detection in all of our in situ samples and must go from bottom water concentration to zero in significantly less than 2 cm (Fig. 4). The boxcore samples show that NO~ goes to zero by about 1.5 cm. A broad maximum in dissolved Mn is seen in both the in situ and boxcore data centered at 15 and 8 cm, respectively. The Si data agree well except that a maximum in the boxcore data at about 5 cm is not seen in the in situ data. At Sta. 183-3, oxygen penetrates to between 5 and 8 cm and NO3 to about 15 cm (Fig. 5). The linear decrease of nitrate through the oxygen-containing layer suggests that nitrate is diffusing through this layer to a deeper denitrification zone. Manganese is low in the top 5 cm and then increases to a broad maximum, with concentrations greater than 40 I~M, centered at 34 cm. The top 50 cm of the sediments at 183-4, 5 and 6 are aerobic red clay (Figs 6-8). At all three locations oxygen decreases but does not reach zero. Oxygen decreases to about 30 l~mol kg-1 at 183-4 and 100 I~mol kg-t at 183-5 and 183-6. Nitrate increases with depth and does not show any decrease which would occur as a result of denitrification. Dissolved Mn was not detected at any of these stations. Dissolved Si increases rapidly to a plateau at all stations and the in situ and boxcore data tend to agree well. The plateau values decrease away from the margin and range from 700 l~mol kg-t at 183-2 to 190 limol kg-t at 183-5 and 183-6. Phosphate was analysed; however the data are not shown here because they were generally less than bottom water values in both the in situ and boxcore sediment samples. We believe that some artifact during sampling acts to remove phosphate from these samples. T T183-STN Oz o 0

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~ zoo o I (=l

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Fig. 5. Interstitial water profiles of oxygen, nitrate, manganese and silica at Sta. 3 of "VF183. The following data are s h o w n : o x y g e n , o n e M sire; nitrate, one in sit,+ and one boxcore; manganese, one in $itu; silica, o n e b o x c o r e and o n e in sit,. See Fig. 4 for explanation of the symbols.

J. W. Mt.~RAY and K. M. K t n v n ~

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Fig. 6. Interstitial water profiles of oxygen, nitrate and silica at Sta. 4 of 1"]'183. The foUowing data are shown: oxygen, two m situ; nitrate, one boxcore and two in situ; silica, two M aim. See Fig. 4 for explanation of the symbols.

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Fig. 7. Interstitial water profiles of oxygen, nitrate and silica at Sta. 5 of "1"1"183. The following data are shown: oxygen, one m $itu; nitrate, one boxcore and o n e / n situ; silica, one boxcore and one in $itu. See Fig. 4 for explanation of the symbols.

Organic matter di~enem

67

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DISCUSSION

Sedimentary organic carbon

EMERSON et al. (1985), EMERSON (1985) and EMERSON and HEDOES (1988) have discussed the factors that control the rate of organic matter degradation and preservation in marine sediments. With regard to the oxygen and organic carbon distributions, two "end member" cases are considered. In the "carbon-limited" case oxygen is not totally consumed and carbon reaches a low but non-zero level representing the refractory organic carbon concentration. The "oxygen-limiting" case applies when oxygen is totally consumed in the interstitial water and residual reactive carbon is buried below the zero oxygen level. Development of a diagenetic model coupling organic carbon and oxygen reactions suggested that the organic carbon content of sediments is most strongly influenced by the flux of particulate organic carbon to the sediments, the bottom water oxygen content and the organic carbon degradation rate constant. Other variables being equal, the concentration of organic carbon in the surface sediments is directly and linearly proportional to the POC flux to the sediments (EMERSONand HEDOES, 1988). Steep gradients have been shown to exist for organic carbon just below the sedimentwater interface, especially when sampled on a mm scale (EMERSONet al., 1985; JARNKEet al., 1986). Even though our sampling interval was only 0.5 cm, we also observed steep interracial gradients at all stations except 183-2. The organic carbon content at 183-2 is higher (by at least a factor of three) and more variable than at the other stations in the transect. Station 183-2 is the closest to the continental margin and because it is located on the Monterey deep sea fan, deposition at this site is probably highly variable (NOR~RK et

68

J.W. MU~A't and K. M. KUIVILA

al., 1984). Moving progressively away from the margin to Stas 183-3, 4, 5 ahd 6, organic carbon decreased by 0.1--0.3% in the top few cm and then decreased more slowly but continuously with depth. Constant background levels of organic carbon at depth were not convincingly observed in the top 20-30 cm of any of these stations. Trace amounts of inorganic carbon were measured at all stations and in some cases (183-6-1,183-5-2, 1834-1) clearly show a systematic decrease with depth. This inorganic carbon is undoubtedly some trace carbonate mineral. Nevertheless, these sediments are for all practical purposes typical red clay deposits. Organic carbon has been measured previously at a mid-Pacific gyre site (MPG-1; 29°N, 156°W) (DRuFFEL et al., 1984) which is very close to our 183-6. At MPG-1 organic carbon decreases from 0.3% at the core top to 0.1% below 20 cm. The inorganic carbon was also very low and decreases from 0.0085 to 0.002% at 20 cm. The average sedimentation rate determined by 23°Th/Z32Th is 0.95 mm ky-l. The mixing depth, on a long-term basis, inferred from the penetration of bomb fallout nuclides, appears to be on the order of 20 cm. The stoichiometry of the decomposing organic matter can be determined from the organic carbon and total nitrogen data. Total nitrogen was measured on the same samples as organic carbon. The values are much lower and the scatter is greater, nevertheless the profiles are similar. Total nitrogen decreased with distance from the continental margin and decreased with depth in the sediments at each location. The relationship between total nitrogen and organic carbon yields information about the stoichiometry of the decomposing organic matter. As shown by GRUNDMANtSand MURRAY (1982), if there is a steady-state input of organic matter with a constant C/N ratio, constant fixed plus exchangeable NH4 and constant non-metabolizable organic carbon, the slope of a linear C/N relationship is equal to the C/N ratio of the organic matter being decomposed. The application of this approach to the red clay sediments in the western equatorial Pacific resulted in a good correlation (r = 0.79) between total nitrogen and organic carbon. The atomic ratio derived from the slope was 9.6 and the intercept on the total nitrogen axis suggested a fixed plus interchangeable NI-14content of 0.013%. Direct measurement revealed this to be mostly fixed nitrogen. The organic carbon/total nitrogen plots for TI'183 show two different relationships (Fig. 9). The data from Stas 183-3, 184-4, 183-5 and 183-6 are strongly correlated (r -0.93) and have a slope of 8.6. This corresponds to a molar ratio of 10.0, which is slightly larger than the C/N ratio of 6-7 for phytoplankton and zooplankton (REDFIELDet al., 1963). The intercept on the total nitrogen axis suggests a value of 0.05% for fixed plus exchangeable NH4, which is a factor of four larger than we observed in the western equatorial Pacific. The organic carbon/total nitrogen data from 183-2 have a steeper slope (r = 0.90) that corresponds to a molar ratio of 15.8. Typical terrestrial vascular plant organic matter tends to have a larger C/N ratio (20-200) than marine organic matter (HEDGES and PARrER, 1976; HEDGES et aL, 1986) so that high ratios at 183-2 suggest that there is a component of terrestrial organic matter at that site. The regional variation in oxygen-limiting vs carbon-limiting conditions can be established by examination of the interstitial water oxygen and sediment organic carbon. The solid organic carbon/dissolved oxygen relationships for Stas 183-4, 5 and 6 are shown in Fig. 10. The data from each station have a highly significant linear correlation (r = 0.78-0.91). The slope of the lines are a function of many variables, including the diffusion coefficient for oxygen, the organic carbon decomposition rate constant, the bioturbation

69

Organic matter d i a g e n e s i s

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0 I-0.1

~

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, "Pr1831

2.0

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OrganicCarbon(%) Fig. 9. Total nitrogen (%) against organic c a r b o n ( % ) for the sediments at Stas 2 - 6 of'1"1"183. T h e regression through the d a t a at Sta. 2 is Y = - 0 . 0 1 4 + 0.14x (r = 0.90) while the d a t a at Stas 3 - 6 is d e s c r i b c d by Y = 0.052 + 0.086x (r = 0.94). O.S

Sin

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OxygenO~M) Fig. 10. Organic carbon (%) in the sediments against interstitial water oxygen at Stas 4--6 of "I"I"183. T h e stations are described by the following regressions. Sta. 4 Y = 0.0052 + 0.0034x r = 0.87 Sta. 5 Y --- --0.59 + 0.0059x r = 0.78 Sta. 6 Y = - 0 . 9 6 + 0.0075x r -- 0.91

coefficient, the stoichiometrie ratio relating oxygen to organic carbon and the porosity (GRur~t)MANISand MURRAY, 1982). In the western equatorial Pacific the slope averaged 3.14 x 10-3. Here the slopes range from 3.4 to 7.5 × 10-3. The intercept of the best fit line for Stas 183-5 and 183-6 is negative, indicating that if all the organic carbon were oxidized there would still be oxygen remaining. These stations clearly fall in the carbonlimited region. The relationship for 183-4 goes through the origin, indicating that there is

70

J.W. MURRAYand K. M. KUtVR~

enough oxygen to decompose all of the organic matter provided all of it is degradable. At Stas 183-2 and 183-3, oxygen goes to zero by 10 cm. Thus, Sta. 1834 appears to represent a good estimate of the transition from oxygen-limited to carbon-limited diagenesis for these sediments along the eastern boundary of the North Pacific.

Stoichiometry of decomposing organic matter At the three carbon-limited stations (183-4, 5 and 6) the interstitial water oxygen and nitrate profles are mirror images of each other. Interstitial nitrate increases while oxygen decreases. At Stas 4-6 there appears to be small nitrate maxima in the top 5 cm. This is especially evident in the boxcore data. It suggests that nitrogen is sometimes preferentially regenerated relative to carbon. This feature may be the result of benthic biological processes although it does not appear in the silica data. Plots of oxygen vs nitrate have been used to calculate the stoichiometry (u) of the decomposing organic matter. In the western equatorial Pacific the value of ¢o~n~o, was 10. l _+ 1.5 (GRUND~ZS and MtmRAY, 1982), which is slightly larger than the classic Redfield ratio of 8.6 (REDW~LDet al., 1963). At Manop site S, J A ~ r ~ et al. (1982b) observed a value of 8.5. For the TI'I83 carbon-limited stations the slope of dO2/dNO3 equals 11.5 (Fig. 11). The correlation coefficient (r) of 0.86 is about the same as that reported for similar data by GRUNDMANlSand MURRAY(1982) and JAHNKEet al. (1982b). Using 1.22 as the ratio of the molecular diffusion coefficients, Do/DNo ,, the calculated value of ao~mo, is 14.0. Assuming a Redfield relation between oxygen and organic carbon, this value of O ~ o , translates into a ~_~ of 12.0. This is in reasonably good agreement with the value of 10.0 inferred from the solid organic carbon/total nitrogen data (Fig. 9).

Rain rate of organic matter The rain rate of organic carbon drives the diagenesis in marine sediments. We chose our station locations to coincide with the locations where the flux of particulate organic

1"1"183Oxygen Vm. Nitrate

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Fig. It. Nitrate vs oxygenin the interstitial water of Stas 4-6 of TT]83. The regression is Y = 52.4 - 0.0868x (r = 0.86).

Organic matter diagenew

71

carbon had been measured as part of the VERTEX program (MAgTn¢ et al., 1987). The VERTEX program deployed free-floating particle interception traps for 13-34 days at several depths at four of the stations for which we have interstitial water data. VERTEX Stas 1, 2 and 4 were occupied in June 1984 while Sta. 5 was occupied in July and August 1983. Our TT183 stations were occupied during September 1984. The VERTEX station numbers and organic carbon flux at 2000 m (their deepest trap) are shown for the respective TT183 station numbers in Table 2. The flux at 2000 m varies by a factor of 9 (0.580-0.063 mol C m -2 y-l) along this transect from 183-2 to 183-6. W ~ and CANUEL(1988) have used organic geochemical analyses of VERTEX trap samples from the eastern tropical North Pacific to argue that sinking particles caught by traps are dominated by zooplankton-derived compounds. A VERTEX time series study was conducted at a site near our Sta. 5 from November 1986 to May 1988 (KNAUERe t al., 1988). During periods of high primary productivity (July-September) the f-ratio estimated from measured nitrogen fluxes tended to go down (0.07); during low productivity (November-April) the f-ratio tended to be higher (0.15). The net result is a relatively small seasonal variation in the nitrogen fux. The carbon fluxes, however, do vary seasonally. Observed changes in the CAN ratio of the flux material imply a seasonal variation in the nitrogen retention in the euphotic zone. These results suggest that seasonal variation in the rain rate of organic carbon may be as large as a factor of two at our sites. Similar seasonal variations of the same magnitude have been observed in other parts of the ocean (e.g. DEUSERand Ross, 1980). We will assume here that the values reported by MARnN et aL (1987) are representative of the seasonal average, thus introducing some uncertainty in our interpretation. MAR'nN et al. (1987) fit the VERTEX data to an equation of the form: (1)

F = Floo(Z/IO0) b,

where Fis the flux at any depth, Ft0o is the flux at 100 m (the log-log intercept), Z is the depth in meters and the exponent b is the log-log slope. We have used this equation to extrapolate the flux to the sediment-water interface. These fluxes also are summarized in Table 2. Because the stations become deeper away from the continental margin, the estimated flux to the sediment-water interface varies by a factor of 15 from 0.360 to 0.023 tool C m-2 y-i. The fluxes at the sediment-water interface tend to be about half of the values at 2000 m. There is justification for extrapolating the flux at 2000 m to the sediment-water interface. SuDs (1980) summarized the early sediment trap data and observed that the flux of organic carbon decreased continuously through the abyssal ocean to the sedimentwater interface. His equation predicts that the flux at 5000 m should be about 2.5 times less than the flux at 2000 m. HosJo (1980) summarized the flux of biogenic matter at four Table 2. The flux of particulate organic carbon at 2000 m as reported by MARTIN et al. (1987). The extrapolatedflux at the sediment-water interface w a s calculated using equations of the form: F ffi Ftoo(z/lOO)b Organic carbon flux (mol m -2 y-t)

TTI83 Sta.

VERTEXSta.

2000 m

Sediment-waterinterface(extrapolated)

2 4 5 6

V-I V-2 V-4 IV-5

0.580 O. 171 O. 113 0.063

0.360 0.080 0.047 0.023

72

J . w . MURRAYand K. M. KUMLA

PARFLUX sites and observed that all all stations except the Panama Basin the flux of organic carbon decreases slowly with depth within the bathypelagic layer. Finally, WALSH et al. (1988a) have recently summarized the organic carbon fluxes at three MANOP sites in the Pacific. The data at sites C and H show a decrease in the organic carbon flux with depth, except for the lower water column where it appears that sediment resuspension can increase the flux (WALSH et al., 1988b). They extrapolated their fluxes to the sediment-water interface using an equation of the form: P = P0 e-k:,

(2)

where P is the extrapolated flux, Po is the measured flux at some depth, k is the reaction rate constant and z is depth. Even though these studies use equations of different form to extrapolate the measured fluxes to the sediment-water interface, there is general consensus that the fluxes do decrease, albeit slowly, within the deep sea. Because of the errors of these extrapolations and the seasonal variation discussed earlier, the fluxes must be considered to have an uncertainty of at least a factor of two. Flux of oxygen The organic carbon and interstitial water oxygen profiles can be modeled to obtain information about the flux of oxygen and the rate constant for organic matter decomposition (BERNER, 1980). We will use the theoretical development and assumptions presented by MURRAYand GRUNDMANIS(1980) and GRUNDMANISand MURRAY(1982). A summary of the equations for organic carbon and dissolved oxygen are given in Table 3. EMERSON et al. (1985) coupled these equations with a flux balance for organic carbon at the sediment-water interface, Fc = K (1 - qb) p (dG/dZ), . o,

(11)

where Fc is particulate flux of carbon to the sediment-water interface and ( d G I d Z ) z , o is the carbon gradient at the sediment-water interface, in order to model the dependency of the organic carbon and oxygen profiles on the rain rate of organic carbon. The oxygen data were fit to equation (8) (in Table 3) by minimizing the variance of the actual data points from the exponential curve. The values of O~, AOz and the values of Bc obtained from the oxygen profiles for each station are summarized in Table 4. The coefficient Bc increases by about a factor of 10 from --0.0235 at Sta. 183-6 to -0.182 cm-t. The values of AO2 systematically increase approaching the continental margin. The flux of oxygen into the sediments can be estimated from these best fit curves and Fick's first law: Flux O2 = -Dovpw qb ( d O z l d Z ) z

= O.

(12)

The oxygen gradient at the sediment-water interface can be estimated using the derivative of equation (8). Thus: Flux 0 2 = - Do:,pw ~b B AO 2.

(13)

This is a lower limit on the flux estimate because the actual gradient close to the sediment-water interface may be steeper (e.g. REIMERS, 1987). We used a value for Do,.pw of 6.0 x 10-6 cm 2 s-t as derived by MURRAY and GRUNDMANtS (1980) and the porosity values in Fig. 2.

73

Organic matter diagenesis

Table 3. Summary of the diagenetic equations used to model the distribution of organic carbon and dissolved oxygen in the sediments. The assumptions made to derive these equations are presented in Mummy and GRUNDMANIS(1980) and GRUNDMANISand MURRAY(1982). They include constant porosity, bioturbation mixing coefficient, sedimentation rate and constant organic carbon degradation rate constant over the depth interval modeled Organic carbon K °~G - OG oz 2 w o-T solution for:

z = 0 z --.-, ~

k~G = O

(3)

G -- Go G--~0 2K

z .

(4)

In the presence of non-metabolizable carbon

GT(Z) = G.,. + G=.o exp (B,z),

(5)

where: .

B¢ ffi

2K

(6)

-

Oxygen 02Oz DO: ~ = ko:G m = ko,G~., o exp ( B ~ )

(7)

solution: O:(z) = O" + AO2 exp

(B,:z),

(8)

where O~ is the constant oxygen concentration at depth and: AO..

ko,

G,.,

p(I

-

¢)u

DO., B~ 12 • 100 -~ if ko2 AOz

= k¢

kcG=.o p(l - ~)a Doz B~ 12" 100- ~"

The factor[p(1 - ~)ul/[12. 100. ~] h necessary to convert organic carbon f r o m % t o moles c m ~ .

Variables

K W

k¢ G

G,,

Z

G= Gm.o Gnm

DO: ko= P Q

(9)

bioturbation coefficient sedimentation rate constant first order rate constant for oxidation of organic carbon organic carbon organic carbon at the sediment-water interface dcpth in sediments mctabolizable organic carbon metabolizable organic carbon at the sediment-water interface non-metabolizable organic carbon molecular diffusion coeffic/ent of oxygen in the sediments constant first order rate constant for oxygen consumption density of solid sediments (2.6 g crn -3) porosity (0.90) stoichiomctric factor relating oxygen to carbon (1.30)

(to)

74

J . w . MURRAYand K. M. KUIVILA

Table 4. The best fit parameters for the lit of the interstitial oxygen data to an equation of the form: O,(z) = O" + AOz exp (B,.z), where AO2/8 the total decrease in oxygen from the interface to the depth where it levels off at some constant value, 0~. The values of B¢ obtained from the oxygen profiles are given as Be (0,.). The curves described by these parameters are shown in Figs 5--8. The rate constant/or organic carbon decomposition (k~) is calculated assuming a mixing coefficient (K) of 150 cmz s-t. Theflttr/8 obtained from the derivative of equation (8) at z = 0 or Flux 0 z = DO, _~bB AOz. The residence time of organic carbon (re)/s estim~'e~d as k: t

O"

AO.

Be(O:)

kc

Flux Oz

T,

Sta.

(ttmol kg-I)

(ttmol I~g-t )

(cm -i)

(ky-z)

(mol m -z y-t)

(y)

183-6 183-5 183.4 183-3

100 105 30 0

64 53 125 154

-0.0235 -0.0487 -0.0570 -0.182

0.083 0.36 0.487 4.97

0.020 0.034 0.099 0.417

12070 2810 2050 201

The diffusive fluxes (Table 4) increase from 0.020 mol 02 m-2 y-t in the red clay sediments north of Hawaii to 0.417 tool m -2 y-t at 183-3 which has suboxic hemipelagic sediments. The average flux calculated using the same approach for nine stations in the western equatorial Pacific was 0.027 mol O2 m-: y-t (MuRrAY and GRUNDMANtS, 1980). SMrrH (1987) recently summarized sediment community oxygen consumption at five locations along a transect from a location north of Hawaii (CNP) to the San Clemente Basin. At his station CNP, which is close to our 183-6 the oxygen consumption decreased from 0.109 mol Oz m-2 y-l in June to 0.027 tool O: m -2 y-i in November. The data suggest that a seasonal variation exists. Our value at Sta. 183-6 of 0.020 mol m-2 y-~ was in September. At Sta. G, which is located about midway between our 183-4 and 5 (0.099 and 0.034 mol m -2 y - t respectively), SMt'm (1987) measured oxygen consumption ranging from 0.036 to 0.126 mol Oz m -2 y-t. Microelcctrode oxygen profiles also have been measured at the CNP site by RI':IMERS(1987). The flux calculated from these profiles (measured in October) is about 0.15 mol Oz m-z y - t which reflects the steeper gradient in the top few mm of the sediments measured by the microelectrode technique. The comparison of our flux estimates with those of SMrrH (1987) and Rt:.tMERS (1987) is reasonably good; however, our values tend to be on the low side of their ranges reflecting the differences in our sampling techniques. At our 183-2 oxygen went to zero very rapidly and we could not calculate a flux. JAliNKEand JACKSON(1987b) recently reported oxygen flux measurements made with a benthic lander at a station very close to our 183-2. Their oxygen consumption rate was about 0.6 mol 02 m-2 y-t. JovcE et al. (1986) have recently calculated that the in situ oxygen consumption rate in the bottom I000 m of the subarctic North Pacific is about 0.0025 ml I-l y-t. If all this oxygen consumption occurs at the sediments it is equivalent to an oxygen consumption rate of 0.1 mol O2 m-2 y-t. It appears that benthic oxygen fluxes are the right magnitude to account for oxygen consumption in the abyssal North Pacific. It has been shown that oxygen is quantitatively the most important electron acceptor for the decomposition of organic matter in pelagic marine sediments (JAttNKE et al., 1982b; BENDER and HEGGIE, 1984). This is also clearly seen in our data from the northeastern Pacific over a range of carbon rain rates. The relationship between the flux or organic carbon to the sediment-water interface (Table 2) and the benthic flux of oxygen is excellent (Fig. 12). The data points tend to fall above the 1-1.3 line suggesting that our estimated carbon rain rates at the sediment-water interface may be a little low. Our data support the argument that most of the rain of organic carbon to the sedimentwater interface is decomposed by oxygen.

Organic matter diagenesis 10 0

75 1:1

TT183 ~

3

E g

o v N

_= U.

10"

/ //.

1O-1o.~

Im

4 II////, r

':i . . . . . . . . . . . . . . . 10"I

10 0

CarbonRainRate(molm-2yr-1) Fig. 12. Thebenthicfluxofoxygencalculatedfromtheinterstitialwaterdataagainstthecarbon rain rate, extrapolated to the sediment-waterinterface. The interstitial data from Sta. 3 is paired with the carbon fluxdata from Sta. 2. An arrow in the figureindicatesthat the fluxdata from Sta. 2 may be an upper limit for Sta. 3. Two solid lines are shown, representing oxygen to organic carbon ratios of l to I and 1.3 to [ (REDFIELDet al., 1963). Recently JAtlNKE and JACKSON(1987a) have observed that bcnthic oxygen fluxes are a strong function of water depth and are significantly larger toward the eastern than the western boundaries. Their curve for the eastern boundary was strongly influenced by high fluxes at a single location at the base of the Patton Escarpment. The benthic oxygen consumption figure of JAIINKE and JACKSON (1987a) is reproduced here with the T1'183 data points included (Fig. 13). Our data give support to this general concept of higher fluxes at the eastern boundary, especially below 4000 m. Organic carbon degradation rate constant The exponential parameter (Bc) used to fit the interstitial water oxygen profiles can be used to estimate the magnitude of the first order organic carbon degradation rate constant (kc). Assuming that bioturbation is more important than sedimentation for the downward flux of solid organic carbon (e.g. EMERSON, 1985), then Bc can be approximated by (kJK) t (equation 6 in Table 3). The biological mixing coefficient (K) has been modeled as a random process analogous to molecular diffusion and the value of the mixing coefficient has been treated as a constant in most previous deep sea work (GRUNDMANISand MURRAY,1982; EMERSON et al., 1985) with typical values around 150 cm 2 ky-t (COCHRAN,1985; JAHNKE et al., 1986). GRUNDMANISand MURRAY (1982) showed that the same value of the organic carbon degradation rate constant could be used to model the interstitial water oxygen, nitrate and solid organic carbon profiles, supporting the use of a value for K of 100 cm 2 ky-t. We now know that the bioturbation coefficient in continental margin regions like the Washington slope and rise (CARPENTER et al., 1987), Panama Basin (ALLER and DEMASTER, 1984), New England slope

76

J . w . MuggAY and K. M. Kunm...A BENTHIC .2

0

•

4

5

6

FLUX (tool C l m 2 1 y r } .4 .6 .8

1.0

e3

tU" ~

Fig. 13. Benthic flux (oxygen consumption as a function of area) as a function of water depth for non-upwelling (eastern) boundaries and upweUing (western) boundaries. The solid lines are from JAtlsgE and JACKSON (1987a). The benthic fluxes calculated for T1"|83 Stas 3--6 are shown. A directly measured benthic flux, at a location close to our Sta. 2, was presented by JAIlNKE and JACKSON (1987b) and is shown for comparison.

al., 1988), and Santa Catalina Basin (SMmlet al., 1986) can be more than 100 times larger. In the New England slope study the mixing coefficient was essentially constant, and not a function of depth, for water depths from about 500-2700 m. in addition, we know that bioturbation varies with depth in the sediments (e.g. ALLER and DEMASTER, 1984), reflecting a particle-selective feeding mechanism (e.g. JAHNKEet al., 1986) and, at least on small time and space scales, is influenced by spatially rare but dynamic events (SMml et al., 1986). Thus, parameterizing K as a random process and assuming a constant value over the depth of our profiles is probably not accurate (BOUDREAU, 1986); however adapting an alternate approach would require more detailed information than we have about the bioturbation process at our stations. Thus, as a first approximation, we have assumed that the bioturbation coefficient at 183-3, 4, 5 and 6 is constant and equal to 150 cm 2 ky-1, in order to facilitate comparison with rate constants calculated by EMERSONand HEDOES (1988). With this assumption the values of kc have been calculated from the values of Bc(Oz) presented in Table 4. The resulting values of k¢ increase from 0.083 ky-t at 183-6 to 4.97 ky-t at 183-3. Following a similar approach, the parameter Bc also can be calculated from the nitrate profiles (GRUNDMANISand MURRAY, 1982). Except for Sta. 6 the agreement between Bc(O2) and Be(NOn) is reasonably good (Table 5). The values of B¢ calculated from the nitrate data tend to be higher than those from oxygen, and we believe this suggests preferential regeneration of nitrate relative to carbon. Thus, here we use the values of Be determined from the oxygen profiles to calculate kc. SMml and BALDWIN(1984) and SMrm (1987) have observed that sediment community oxygen consumption varies seasonally in the central and eastern North Pacific. Presum(ANDERSON et

77

Organic matter diagcnesis

Table 5. co.wa,~on of values of Be obtained ffom the 02 and NO3 profiles. The lines described by both parameters are drawn through the nitrate data in Figs 6--8

Sta.

B~(O2)

Be(NO3)

183-6 183-5 183-4

--O.O235 -0.0487 -0.0570

0.099 0.0536 0.0820

ably this reflects seasonal variations in the rain rate of organic carbon. Modeling studies by MARTINand BENDER(1988) argue that such seasonality should not be reflected by the benthic oxygen fluxes unless the organic carbon degradation rate constant is on the order of 4 y-l. This value is about 100 times greater than our estimates of kc; thus seasonal variability is probably not an important factor for the data. Nevertheless, an implicit assumption in the calculations presented here is that the values of the rain rate of organic carbon (Table 2) approximate the annual average rain rate. Both the rain rate of organic carbon to the sediment-water interface and the organic carbon degradation rate constant increase along our transect from the central North Pacific to the continental margin. The range of organic carbon rain rates is about 50. Over this range the degradation rate constant varies by approximately 10s and the loglog relationship is highly linear. These results support the argument that at constant bottom water oxygen concentrations the flux of carbon is a master variable for the degradation rate constant in marine sediments. Residence time of organic carbon The reciprocal of the organic carbon degradation rate constant is the residence time of organic carbon with respect to oxidation in the upper layers of the sediments. These values range from 12,000 years at 183-6 to 200 years at 183-3. For comparison, the corresponding valucs determined in the western equatorial Pacific average 250 years (GRUNDMANIS and MURRAY, 1982) and 15-100 years at MANOP sites M, H, C and S (EMERSON et al., 1985). EMERSONet al. (1987) have recently used a 14C dating approach to show that significant fractions of the organic carbon in the surface sediments of the continental margin and equatorial Pacific are degradable on times scales less than 1000 years. They also determined the 14C age of the organic carbon at Sta. MPG-1 of DRUFFELet al. (1984). The residence time calculated for the degradable organic carbon was greater than 10,000 years. In this case the residence times of degradable organic carbon and sediment in the mixed layer (-,-12,000 years) are about the same. When sedimentation rates are as low as they are at this site (~0.1 cm ky-I) and kc is less than 10-4 y - I sedimentation will determine the organic carbon content of the mixed layer (EMERSONand HEDGES, 1988). The decrease in residence time we observe going from 183-6 to 183-3 is consistent with 14C measurements at other locations. CONCLUSIONS

The flux of organic carbon to the sediment-water interface and the organic carbon content of the sediments decrease from the continental margin off California to the middle of the central North Pacific gyre. Station "1~183-4 appears to present the boundary between aerobic and suboxic diagenesis along this transect.

78

J . w . MURRAYand K. M. KUIVlLA 10,000

1,000 I 100

SITE C

-I

[]

I

i 10

CALIFORNIA BORDERLAND

I w. ATLANTIC SLOPE

T

tJ .a¢

SITE l.,.~,.e STN 3 M

l

t eSTN 4 STN 5 • O.1

O.01

I I

.01

S+N s

I

0.1

I

t

I

10

log Fc (mol m-2 y-l)

Fig. 14. Log of the organic carbon degradation rate constant against the log of the rain rate of organic carbon to the scdimcnt-water interface. The values calculated here for TI'183 Stas 3--6 are compared with values from MANOP sites M, tl and C, the Western Atlantic slope and the California Borderland presented by EMEaSON (1985) and EMtmSON et al. (1985). We assumed that the organic carbon rain rate from 183-2 holds for 183-3 and use an arrow to indicate that the actual flux may be smaller.

The flux of oxygen into the sediments agrees well with the organic carbon rain suggesting that most of the organic carbon is decomposed by oxygen within the sediments. The first order degradation rate constant for organic carbon can be estimated from the porewater oxygen profiles. The values increase from 0.083 to 4.97 ky-t moving toward the continental margin. These results show that this rate constant is a function of the organic carbon rain rate. The residence time of organic carbon with respect to oxidation ranges from about 12,000 years in the central gyre to 200 years at the ocean boundary. Acknowledgements--These samples were collected and analysed with the assistance of H. Jannasch, C. Stump, B. Honeyman, B. Wallin, B. Paul and the captain and crew of the R.V.T.G. Thompson. L. Lu conducted about half of the organic carbon, total nitrogen and total carbon analyses. Univcrsity of Washington reviews were provided by S. Emerson and J. Hedges. The comments by W. Martin and R. Jahnke benefited the manu~ript. The manuscript was typed by S. Dow. The research was supportcd by NSF Grant OCE 8416258. University of Washington Contribution no. 1843. REFERENCES ALLF.R R. C. and D. J. DE',tASTER (1984) Estimates of particle flux and rcworking at the dccp-sca floor using :2"~I'h/z3HUdisequilibria. Earth and Planetary Science Letters, 67, 308-318.

Organic matter diagenesis

79

ANDERSON R. F., R. F. Boer, K. O. BUESSELER and P. E. BtsCAYE (1988) Mixing particles and organic constituents in sediments from the continental shelf and slope off Cape Cod: SEEPml results. Continental Shelf Research, 8, 925.-946. BENDER M. L. and D. T. HEC,GIE (1984) Fate of organic carbon reaching the deep sea floor: a status report. Geochimica et Cosmochimica Acta, 48, 977-986. BERNER R. A. (1980) Early diagenesis: a theoretical approach. Princeton University Press, Princeton, NJ, 250 pp. BOUDREAU B. P. (1986) Mathematics of tracer mixing in sediments. II. Nonlocal mixing and biological conveyor-belt phenomena. American Journal of Science, 286, 199-238. BREWER P. G. and D. W. SPENCER (1974) Colorimetric determination of manganese in anoxic waters. Limnology and Oceanography, 16, 107-110. BROECKERW. S., D. W. SPENCERand H. CRAIG (1982) GEOSECS Pacific Expedition, Vol. 3, Hydrographic Data. National Science Foundation, Washington, D.C. CARPENTERR., T. M. BEASLEV,D. ZAHNLEand B. L. K. SOMAYAJULU(1987) Cycling of fallout (Pu, 24tAm, I37Cs) and natural (U, Th, 21°Pb) radionucfides in Washington continental slope sediments. Geochimica el Cosmochimica Acta, 51, 1897-1921. COCHRAN J. K. (1985) Particle mixing rates from MANOP sites in the eastern Equatorial Pacific: evidence from zwPb, 2~'2~Pu and t32Cs distributions. Geochimica et Cosmochimica Acta, 49, 1195-1210. DEUSER W. G. and E. H. Ross (1980) Seasonal change in the flux of organic carbon to the deep Sargasso Sea. Nature, 283, 304-305. DRUFFELE. R. M., P. M. W~LIAMS, H. D. L~NGSTON and M. KtODE (1984) Variability of natural and bombproduced radionuclide distributions in abyssal red clay sediments. Earth and Planetary Science Letters, 71, 205-214. EMERSON S. (1985) Organic carbon preservation in marine sediments. In: The carbon cycle and atmospheric CO2: Natu.'al variations Archean to present, E. T. SUNDQUIS'rand W. S. BROECKER,editors, Geophysical Monograph, 32, AGU, Washington, D.C., pp. 78--87. EMERSON S. and J. i. HEDGES (1988) Interpreting the organic carbon content of marine sediments. Paleoceanography, 3, 62 !-634. EMERSON S., R. JAtINKE, M. BENDER, P. FROELICll, G. KLINKIIAMMER,C. BOWSERand G. SETt.OCK(1980) Early diagenesis in sediments from the eastern equatorial Pacific. 1. Porewatcr nutrient and carbonate results. Earth and Planetary Science Letters, 49, 57-80. EMERSON S., V. GRUNDMANISand D. GRAltAM (1982) Carbonate chemistry in marine porewatcrs: MANOP sites C and S. Earth and Planetary Science Letters, 61,220--2.32. EMERSON S., K. FISOIER, C. REIMERSand D. th:.G(;IE (1985) Organic carbon dynamics and preservation in deep sea sediments. Deep-Sea Research, 32, 1-21. EMt.:RSONS., C. SLUMP, P. M. GR(~OTI.~S,M. S'IUIVER,G. W. FARWEt.t.and F. i I. S¢'IIMIDT(1987) Estimates of degradable organic carbon in deep-sea surface sediments from 14C concentrations. Nature, 329, 51-53. GRUNDMANISV. and J. W. MtmRAY (1982) Aerobic respiration in pelagic marine sediments. Geochimica et Cosmochimica Acta, 46, 1101-1120. llEDC;ES J. I. and P. L. PARKER (1976) Land-derived organic matter in surface sediments from the Gulf of Mexico. Geochimica et Cosmochimica Acta, 40, 1019-1029. IIEI)GES J. i. and J. H. STERN (1984) Carbon and nitrogen determinations of carbonate-containing solids. Limnology and Oceanography, 29, 657. llEDc;I.'.S J. l., W. A. CLARK, P. D. OUAY, J. E. RIOIEY, A. tl. Dt~VOL and U. deM. SANTOS (1986) Compositions and fluxes of particulate organic material in the Amazon River. Limnology and Oceanography, 31,717-738. tloNJo S. (1980) Material fluxes and modes of sedimentation in the mcsopclagic and bathypelagic zones. Journal of Marine Research, 38, 53--97. lloNJO S., S. J. MANGANINIand J. J. COLE (1982) Sedimentation of biogenic matter in the deep ocean. DeepSea Research, 29, 609-625. JAHNKE R. A. and G. A. JACKSON(1987a) The role of the seafloor in maintaining deep-ocean chemistry. Nature, 329, 621-623. JAIINKE R. A. and G. A. JACKSON(1987b) Benthic fluxes on a basin scale: Importance of the eastern boundary. EOS, 68, 1748. JAIINgE R. A., S. R. EMERSONand J. W. MURRAY(1982a) A model of oxygen reduction, denitrification and organic matter mineralization in marine sediments. Limnoiogy and Oceanography, 27, 610--623. JAIINKE R., D. HEGGIE, S. EMERSONand V. GRUNDMANIS(1982b) Porewaters of the central Pacific Ocean: Nutrient results. Earth and Planetary Science Letters, 61,233--256. JAtINKE R. A., S. R. EMERSON,J. K. COCHRANand D. J. HtRSCHUERO0986) Fine scale distributions of porosity and particulate excess 21°Pb, organic carbon and CaCO3 in surface sediments of the deep equatorial Pacific. Earth and Planetary Science Letters, 77, 59-69. JOVCET. M., B. A. WARRENand L. D. TALLEY(1986) The geothermal heating of the aby~al subarctic Pacific Ocean. Deep-Sea Research, 33, 1003-1015.

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