Isotopic fractionations during oxygen consumption and carbonate dissolution within the North Atlantic Deep Water

Earth and Planetary Science Letters, 49 (1980) 485-498 © Elsevier Scientific Publishing Company, Amsterdam - Printed in The Netherlands

485

[61

ISOTOPIC FRACTIONATIONS DURING OXYGEN CONSUMPTION AND CARBONATE DISSOLUTION WITHIN THE NORTH ATLANTIC DEEP WATER

P. KROOPNICK Department of Oceanography and Hawaii Institute of Geophysics, Honolulu, HI 96822 (U.S.A.]

Received February 15, 1979 Revised version received August 30, 1979

The trajectory of the North Atlantic Deep Water is traced from 65°N to 20°N latitude. Along this track the dissolved 02 decreases, the 6180 of the dissolved 02 increases, and the 14C content of the water decreases. From these observations the rate of in-situ 02 utilization in the deep water is calculated to be 0.10 ~mol kg -1 yr -1 . As was observed previously in the Pacific, the 618 0 data presented here indicate that the utilization is probably caused by bacterial respiration. Carbon dioxide is being added to the water at the rate of 0.07 ~mol kg -1 yr -I from the oxidation of this organic matter. An additional 0.12 ~mol kg-1 yr -1 of CO2 is derived from the dissolution of particles of CaCO3.

1. Introduction Variations in the concentrations of dissolved 02 and carbon in the deep ocean below 1 km are controlled by the horizontal and vertical mixing of waters, the rate of input of particulate material from the surface, the rate o f the in-situ process of bacterial respiration and for carbon the rate of dissolution of carbonate tests. The importance o f each of these effects varies with location and must be assessed for each major ocean basin and water mass before a description o f the dynamics of the carbon and oxygen cycles can be completed. Oxygen isotopes are fractionated during respiration, the light isotope 160 being consumed more rapidly than the heavy isotope 180. The degree o f this fractionation and the gross changes in concentration can be used to study the mechanism and rate o f in-situ 02 utilization in the deep sea [1,2]. The concentration of total inorganic dissolved CO2 (ZCO2) is related to both the rate of oxidation of organic matter and the rate o f dissolution of the skeletal parts of foraminifera, coccolithophora, and other CaCO3HIG Contribution No. 975. GEOSECS Publication No. 120.

secreting organisms. Changes in the alkalinity and the 813C of the ~CO2 can be used to assess the relative proportion of CO2 added from respiration and from carbonate dissolution [ 3 - 5 ] . Finally, the changes in the concentration of the isotope 14C caused by radioactive decay can be used to calculate the mixing time for the oceans and the absolute rates for the processes of O: utilization and CO2 production [ 3 , 6 - 8 ] . This paper summarizes and integrates the results from dissolved 02, 8180 of the dissolved 02, ~CO2, 813C o f the ECO2, and the M4C data obtained during the 1972 Atlantic GEOSECS expedition. The data for the above parameters have been presented elsewhere [ 9 - 1 3 ] except for the 6180 results, which are presented here for the first time. The hydrography of the Atlantic Ocean is dominated by the horizontal advection and vertical mixing of water mass cores formed in high latitudes. A vertical section thus intersects several layers, each differing in chemical properties and place of origin. For this reason, vertical diffusion-advection models used with great success in the Pacific Ocean [3,14] cannot be applied in the Atlantic, and we must be content with the less sophisticated model presented here. This model considers the North Atlantic Deep Water (NADW) as a partially closed system in which

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487 there are (1) particulate input from falling detritus, (2) no other inputs or outputs, and (3) in-situ utilization and production of dissolved 02 and CO2, respectively.

2. Experimental methods Samples were collected between July 1972 and March 1973 on board the R/V "Knorr" during the Atlantic GEOSECS expedition. The ship track and the position of each sampling site are shown in Fig. 1. Isotopic analysis of the dissolved 02 requires that the 02 gas be quantitatively extracted from 1 liter of seawater. Previous methods for the shipboard extraction of such a large volume of dissolved gas involved either boiling the seawater sample in a vacuum while pumping continuously with a toepler pump to transfer the gases into glass flasks, or using a condensible carrier gas such as CO2 to strip the dissolved 02 from the seawater, followed by adsorption onto molecular sieve [2]. A more convenient method was required for the GEOSECS expedition, since various technicians had to operate the extraction system under adverse conditions over a 9-month period. The method adopted involved equilibrating the sample of seawater with an evacuated vo!ume [15] followed by adsorption of the equilibrium mixture of N2, 02, CO2, water vapor, and minor amounts of other gases on molecular sieve. Seawater samples were collected with 304iter PVC Niskin bottles to insure that no reaction of 02 with metal occurred. Concentrations of dissolve d 02 and 2;CO2 were measured at sea on the same samples [16]. The water samples were transferred immediately to 1-liter glass reagent bottles, which were filled from the bottom and allowed to overflow for at least half of their volume to exclude air. Saturated HgCI2 solution (0.5 ml) was added to suppress bacterial growth, but most of the samples were extracted within 24 hours of collection. Fig. 2 shows the portable tandem vacuum system used for the large-volume total gas extraction. The reagent bottle was attached via the siphon tube and air was flushed from the line through the three-way valve. The water was forced through a removable glass flit into the previously evacuated extraction flask. At modest flow rates equilibration of the gas between the water and the air space occurred instantaneously. When the lower bulb

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was full, the flow of water was stopped and the extracted gases were isolated by closing the stopcock between the two bulbs. The accumulated gases were then adsorbed at -183°C by type 5A molecular sieve placed in the bottom of the evacuated breakseal tube. The extraction efficiency was monitored by collecting the water from the lower bulb and comparing its dissolved 02 content with that in the original sample; 92-98% of the dissolved 02 was extracted in this manner. The equilibrium isotope fractionation between the dissolved and gaseous 02 is known [17], and a mass balance calculation indicates that for an extraction yield of 95%, a correction of only 0.04°/oo must be added to the measured 5180 value. In the laboratory, CO2 and any water present in the breakseal tube are separated by passage through two U-traps at -196°C. Oxygen is then converted to CO2 for isotopic analysis by combustion on a carbon rod, according to a combustion procedure described previously [2]. The CO2 samples are analyzed on a Nuclide 3"-60 ° or 6"-60° double collection mass spectrometer [18]. The data are corrected for instrumental effects and mass spectral cross terms [19]. Results are reported in the delta nomenclature relative to atmospheric 02 : 6180 = [(R/Rstd)- 1] X 10 a where R = la0/160 and Rst d refers to atmospheric 02. To refer the data to the SMOW isotopic standard, the value of the fractionation factor between CO~

488 and water is required. Kroopnick and Craig [2] discuss this problem in detail and give the necessary equations. Due to uncertainties in sample collection, extraction at sea, and combustion to CO2, the precision for the dissolved 02 8180 data is +0.1 °/oo. However, individual samples within a profile frequently exhibit variations much greater than 0.1 °/oo due either to the leakage of air into the samples causing a decrease in the 6180, or to inadequate poisoning upon collection, which would allow respiration to continue in the bottle and cause an increase in the measured 8180.

3. Results

3.1. Oxygen isotopes The 6180 profiles for the GEOSECS Atlantic stations are shown in Figs. 3 - 5 *. The results are reported versus air, thus the surface samples have 8 a80 values near zero, indicating that the water has equilibrated with the atmosphere [17]. Values below zero represent samples in which the rate of photosynthesis was greater than the rate of respiration. Due to the layered structure of the Atlantic Ocean, the individual profiles of dissolved 02 and 8180 do not show the clear maximum at the 02 minimum followed by a gentle decrease with depth, as was seen in the Pacific samples [2]. Instead, many of the profiles are virtually constant with depth below the shallow 02 minimum and 6180 maximum. The GEOSECS II profile was described in detail by Kroopnick et al. [20]. In general, the deep water 8180 values for the western Atlantic (Fig. 3) are enriched about 40/00 relative to atmospheric 02. In the central Atlantic (stations G II to 58), where a pronounced 02 minimum occurs between 500 and 1000 m [9], enrichments of up to 6 % 0 are observed. The nutrient-rich Antarctic Bottom Water (AABW) and Antarctic Intermediate Water (AAlW) show large dissolved 02 depletions and corresponding 8180 enrichments. As these

* The 02 isotope data are on file with the GEOSECSOperations Group, Scripps Institution of Oceanography, La Jolla, California, and at the National Oceanographic Data Center, Washington, D.C.

enriched waters spread northward and are diluted by the less enriched NADW, the Atlantic Basin as a whole assumes characteristic 8180 values of about 40/00 in the south and 3.5°/oo in the north. At higher latitudes, such as stations 19, 17, 11,5,69, and 74, the water column is almost completely mixed vertically. The most southerly station (78) is anomalous and will not be discussed further. Samples from stations 89-111 in the southeast Atlantic (Fig. 4) show very high enrichments of almost 8o/00 near the shallow dissolved 02 minimum, as is expected for low-O2 water beneath a highly productive area. The dissolved 02 concentration in the 02 minimum for stations 89-111 ranges from 186 to as low as 84/~mol/kg, indicating a large depletion from the 300 ~tmol O2/kg expected if the water were in equilibrium with the atmosphere at this temperature. The samples from the northeast Atlantic (Fig. 5) show enrichments which are much less pronounced due to the lower phytoplankton productivity in this area.

3.2. Variations within the core o f the N A D W In order to study the processes occurring within the NADW, the core of this water must first be identified. Water mass formation occurs in both the Norwegian and Greenland Seas by the cooling of Atlantic Ocean surface waters flowing northward through passages around Iceland. The newly formed high-density, high-salinity, low-nutrient [21] waters are contained in deep basins with shallow sills. The overflows from these basins follow complex trajectories but eventually combine with a water mass produced in the Labrador Sea to produce a composite water mass known as the NADW. In the central Atlantic the NADW becomes sandwiched between northward-flowing AAIW and AABW which have lower salinities and higher nutrient concentrations [9.]. Detailed study of the GEOSECS temperature, salinity, density, and especially dissolved 02 data [22] indicates that the NADW can be traced from 65°N to nearly 30°S. The NADW is identifiable over these latitudes as a salinity maximum of about 34.94 -+0.01°/oo and a sigma theta of 27.90 -+ 0.05. The dissolved 02 decreases southward but vertical profiles show constant values over a range of 1 2 5 400 m above and below the salinity maximum, indi-

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cating a homogeneous body of water at each position. This core of NADW is graphically shown in Fig. 6 where the cross-hatched area represents the depths from which the samples discussed in this section were collected. To study the chemical processes occurring within the NADW the GEOSECS dissolved 02, 8180 of dissolved 02, ECO2,613C of the ECO2, alkalinity and 6 I4C data were averaged over the depth range that includes the core of the NADW (Fig. 6). The dissolved 02, ECO2, and alkalinity data are taken from the GEOSECS leg reports [9,22]. 618 O is reported here for the first time. The 8 I4C data are from Stuiver [12,13] and Ostlund et al. [11]. 813C data are reported by Kroopnick [10]. The ZCO2 data published in the leg reports have been modified as sug-

gested by Takahashi et al. [23] ; 16/lmol/kg has been added to data from stations 3,5, and 11 since they were measured by gas chromatography in the laboratory, whereas all the others were measured by potentiometric titration on board the ship. The dissolved 02 decreases by 31 bunol/kg between 65°N and 20°N (Fig. 7). This decrease is almost linear and corresponds to an increase in the ECO2 of 58/~mol/kg over the same range of latitude. The 614C data decrease over the same region by 500/00. Stuiver [13] plotted 814C and dissolved 02 over these same latitudes; he observed overall decreases but erratic trends because he defined the NADW core solely on the basis of a salinity of 34.9060/00. Identifying the core of the NADW by

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salinity, density, and dissolved 02 (Fig. 6) as was done here produces much more linear a4C and dissolved 02 curves. The GEOSECS station track (Fig. 1) indicates that from stations 11-26 samples were taken from the relatively narrow area of the North Atlantic. At station 27 the western Atlantic basin opens into a broad abyssal plain. At station 36 (15°N)the track swings eastward following the coast of South America and causes the sudden change in the slope of the data at 20°N. South of station 36 the GEOSECS expedition no longer accurately sampled the core of the NADW. Between 20°N and 30°S dissolved 02, Y-CO2, and 14C all exhibit large fluctuations but the trends are similar.

The decrease in dissolved 02 of 31 ~tmol/kg between 65 ° and 20°N is accompanied by an increase in its 6a80 of 1.4O/o0 (Fig. 7). In order to explain this enrichment I shall assume that the southward-flowing NADW can be treated as a closed system. It follows that any changes in the chemical parameter can be caused only by in-situ processes. If the observed rates of 02 consumption, ~CO2 production, and 61 s O fractionation along the horizontal plane of the NADW are similar to the rates observed in the Pacific Ocean where an open system model can be applied, we would presume that similar biological processes are operating. On the other hand, if radically different results were obtained we must assume that the nature of these in-situ processes has not yet been characterized correctly. In closed-system respiration experiments, surfacedwelling bacteria fractionate 180 relative to 160 by a factor of 21 °/oo [1]. In the deep sea, however, application of the open system vertical diffusionadvection model yielded an eflrichment factor of only about 10°/0o [2]. This difference was attributed to the colder water temperatures and to the different species of bacteria found at these deeper depths. The closed-system Rayteigh distillation model will now be used to calculate the fractionation of 180 during the consumption of dissolved 02 within the NADW [1 ]. This fractionation is usually expressed as the enrichment (e) calculated using the equation:

R/Ro =f-e where R is the ratio 1~0/160 of the dissolved 02 being removed at any instant, Ro is the ratio of the remaining 02, and f i s the fraction of dissolved 02

492 remaining. Since fi 180 increases as the O2 concentration decreases, the enrichment factor is positive. A more convenient form of this equation which uses the tabulated 8 value is: X - ?to = - e - I n ( f ) where ?t = ln(1 + ~5180). Substituting the changes in dissolved 02 and 6180 observed in NADW into the above equation yields an enrichment of 130/00. This is quite similar with the 10°/0o enrichment found previously [2] and suggests that much the same metabolic processes are operating within the core of the southward flowing NADW and the water column of the eastern tropical Pacific Ocean.

4.2. Carbon The 14C data can be used to calculate the rate of 02 consumption. However, the assumption of a closed system must be modified to allow the addition of CO2 from two sources. The first is from the oxidation of organic matter which increases the ZCO2 in direct molar proportion to the amount of dissolved 02 used. The second source is the dissolution of the CaCO3 of skeletal parts of organisms that live in the overlying waters but sink to the bottom after death. The change in alkalinity can be used to measure the amount of CaCO3 dissolved in the water, since for each/lmole of CaCO3 that dissolves, the alkalinity increases by 2/aequivalents and the ECO2 increases by 1 ~tmole. Thus the proportion of inorganically produced CO2 is: F c o 3 = ~ Aalk/AECO2 The alkalinity data for the extreme North Atlantic are not available but extrapolation of the data from stations 33 through 27 to station 11 indicates a net change of about 75 ~tequivalents/kg. The ECO2 change between 65 ° and 20°N is 58/amol/kg (Fig. 7). The inorganically derived fraction of ECO2 is thus about 65% of the total. Performing this calculation only between 20 ° and 40°N where both alkalinity and ECO2 data are available, one finds that F c o s = 90%. These values are considerably higher than previous estimates for the whole Atlantic of 20-30% [4] and also for global estimates that considered both the Atlantic and Pacific Oceans [5,24]. The North Atlantic generally is an area of relatively low produc-

tivity and the NADW has a relatively low organic carbon content. Thus, it is to be expected that the percentage of CO2 produced by the dissolution of CaCO3 appears higher than average. The 61 aC of the ECO2 data can also be used to test this model of the relative sources of the CO2. The average 613C of the CO2 being added can be calculated from the mass balance equation: ( ~ C O 2 • 613C)initial + ( ~ C O 2 • /513C)aaded = (Y, CO 2 . (~ 13C)final

Substituting in the values of 2154 and 2212 for the Y,CO2 at 65 ° and 20°N, and 1.0 for the 813C values yields the result that the 613C of the added CO2 is about 0.0°/oo. The C02 produced from the dissolution of biogenic CaCOa has a 8 laC of +20/00 while CO2 from oxidation of marine organic matter has a 8laC of -230/00 thus the laC data imply that 92% of all the CO2 added to the NADW between 65 ° and 20°N is derived from carbonates. While the two estimates of 90% are much too high, the discrepancy between the data presented here and previous data [4] cannot be reconciled. The maximum possible change in 61 aC is 0.1 °/oo which would lower the carbonate fraction to 80% (the 0.1°/oo offset between UH and SIO data discussed previously [10] does not apply since all tile SIO calculations are south of 15°N and the calculations presented here use only UH data).

4.3. Redfield ratio value The Redfield ratio or RKR value provided a final check on the consistency of the COz and 02 data. The proportion in which the elements of seawater enter the biochemical cycle is determined by the composition of the biomass. Redfield et al. [25] summarized the available data and found that during the oxidation of plankton, 138 moles of 02 are consumed for every 106 moles of carbon oxidized:/xO2 : ACO2 = 1.3. The fraction of ECO2 produced from the oxidation organic carbon is then: AO2 F°rg" - RKR • AECO2 Vertical diffusion-advection and other models for the eastern tropical Pacific gave RKR values from 1.20 to 1.45 [5,26], and the value of 1.3 is often

493 accepted as the metabolic factor for the fixation and oxidation of organic matter. Substituting 1.3 for RKR, and 31 and 58/~moles for AO2 and A~;CO2, respectively, in the above equation, I calculate that Forg. = 41% (Fco3 = 59%). The observed changes in ECO2 and dissolved O2 are 58 and 31 ttmoles, respectively. To explain the discrepancy noted above the 2CO2 change would have to be 4 times larger than observed; clearly this is absurd. Thus, the reason for this much too large apparent carbonate fraction cannot be reconciled at this time, but when considered together the closeness of the above estimates indicate that the predominant flux of CO2 into the NADW comes from the dissolution of falling particles of CaCO3.

4.4. Aging So far nothing has been said of the absolute rates for any of these processes but only of their relative effects. Stuiver [13] showed that the Aa4C of the NADW decreases as it flows southward and used this decrease to calculate an age. The first and predominant effect is due to the aging or decay of the radioactive fraction of the 2;CQ. Every 83 years the Ax4C of the ~CO2 will decrease by 10U/oo. The addition of CO2 from CaCO3 dissolution and from the oxidation of organic matter, as discussed previously, can modify the apparent 14C activity. The 14C activity of the dissolved organic carbon in the Pacific is very low [27], whereas the activity of calcareous matter also collected in the deep Pacific is high and is similar to that observed in surface waters [28]. To correct for the addition of C02 from these sources, we shall assume that, as in the Pacific, the added organic carbon has a A14C of -3500/00 [27] (maximum age), and that the A14C of the carbonate carbon is +100% o as measured for surface water in this area [12]. If 35% of the added carbon comes from organic matter then the A14C of the 58 pmol ECO2/kg added from both sources would be -60O/o0. Due to the large amount of 14C produced by nuclear weapons since 1945, it is difficult to assess the A~4C level of the NADW at the time of its formation, but it was probably close to -800/00 [13]. If the 58//moles of detrital C02 per kg with a A14C of +100% o were added since 1954 to the 2154 gmol CO2/kg present when the NADW left the surface, the

14C activity would increase by 5o/00, causing the seawater to appear about 40 years too young. Clearly~ this is the maximum offset one could expect. The actual error is probably on the order of only four years. On the other hand, if we were to assume instead that the dissolving organic matter has incorporated bomb-produced carbon with a A14C of +100°/oo [29] instead of the -350°/oo value used above for the residual organic matter [27], then the average activity of the added CO2 would again be nearly +100°/oo, suggesting that a 40-year offset is possible. The age of the NADW can now be computed from the radioactive decay of 14C. For calculation, the absolute concentration of ~4C must be used, since the A14Ca- value represents the "tabulated" 14C/a2C ratio ot a sample relative to an arbitrary standard. The tabulated values were normalized for laboratory isotope fractionation by measurement of the 613 C in the final sample counted for 14C. This procedure fails to take into account the actual natural variations of 13C in the ocean environment [10]. Following the notation adopted by Craig [14], the A14CT values are back-corrected to the actual in-situ ~4C concentrations by means of the following algorithm: A14C A

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in which the subscript "A" denotes the actual X4C and aaC values in the original water sample. The factor 2 outside the parenthesis denotes that 14C fractionation effects are twice the effects for ~3C. The value 2 inside the parenthesis is the mean value for 6 ~3C of surface ocean water. The "absolute" ~4C concentrations are obtained from the measured ~CO2 concentration (C) by: 14C = C(1 + 10 -3 A14CA) . The GEOSECS A14C data [11-13] have been recalculated using the above equations and the absolute 14C concentration is plotted in Fig. 7. The corrections are small (~17 years), but are included for completeness. The resulting plot is linear and indicates an apparent aging for the NADW of 304 years between 65 ° and 20°N (350 years if the maximum corrections for the addition of biological and detrital bomb carbon are used). This age is about twice that given by Stuiver [13]. As discussed above the principal difference in the two calculations is that Stuiver defined the core of the

494 NADW differently (a small correction for laC effects on Aa4CT has also been included here). Extension of this model into the Southern Hemisphere is unreliable because extensive mixing occurs between the AABW and NADW at these latitudes. Assuming the AABW is substantially older than the NADW, any calculated age would necessarily be a maximum age.

4.5. Rates o f oxidation and dissolution The in-situ rate of 02 utilization in the NADW can now be evaluated from the net change in 02 calculated earlier (31/2mol/kg) and the elapsed time of 304 years. The resulting rate of 0.10/2mol kg -~ yr -1 is in excellent agreement with the rates of 0 . 0 8 0.17/amol kg -1 yr -1 calculated for the eastern equatorial Pacific with the vertical diffusion-advection model [2,5]. A more detailed numerical model for the deep Pacific Ocean [8], in which both the horizontal and vertical distribution of dissolved 02 were considered, also gives an average 02 consumption rate of 0.10/amol kg -1 yr -1 for the deep water of the Pacific between 1 and 5 km. The addition of CO: to the water column similarly

is calculated to be 0.19 ~tmol kg -1 yr -1. Since about 35% of this addition is due to the oxidation of organic matter, I conclude that CaCO3 is dissolving at the rate of 0.12/amol kg -1 yr -1 . Over a 4-kin water column the flux would be 4 8 / l m o l cm -2 yr -1 . If the rate of accumulation of CaCO3 in the sediment beneath the Sargasso Sea is taken as 1 ram/10 a yr (2/lmol cm -2 y r - l ) , then about 96% of the sinking CaCO3 particles dissolve in the water column. This result for the NADW can be compared with the integrated fluxes of CaCO3 estimated by Krishnaswami et al. [30], which were derived by comparing the production and deposition rates of CaCO3 for several areas of the Atlantic Ocean. Recalculating their results to give the average dissolution over a 4-km depth interval, we find values ranging from a high of 0.16/amol kg -1 yr -1 for the Argentine Basin to a low of 0.03 for the equatorial Atlantic (Table 1). Lerman and Lal [31] developed a theoretical model to calculate the regeneration rate of carbonate based on the abundance of particles in the biologically productive zone of the ocean, the shape of the particle-size distribution curve, and the dissolution rates of calcite and aragonite. Their revised estimate is that 12 +

TABLE 1 Summary of utilization and production rates * Property 02

CO2 CaCOa

Rate (vmol kg-1 yr -1 ) -0.10 -0.12 -0.12 -+ 0.04 - 0.10 0.19 0.11 +- 0.07 -0.12 -0.06 -0.35 -0.16 -0.09 -0.02 -0.04

Location

Method of calculation **

NADW Atlantic Pacific Pacific

this work change in 02 along density surface [6,33] vertical diffusion-advection model [ 2,5 ] horizontal-vertical model [ 8]

NADW Pacific

this work vertical diffusion-advection model [ 2,5 ]

NADW world ocean Cape B a s i n Argentine B a s i n Canary B a s i n equatorial Atlantic world ocean

this work box model [4] production-deposition [30] production-deposition [30] production-deposition [ 30] production-deposition [30] particle size distribution [ 31,32 ]

* Average value for the depth interval between 1 and 5 km. ** Numbers in brackets refer to references listed at the end of this paper.

495 6/lmol cm-2 yr-1 of CaC03 is regenerated between 1 and 5 km, which is equivalent to 0.03/1tool kg -1 yr -1 [32]. Using a procedure similar to that employed in this work, Li et al. [4] obtained a result of 0.06/amol kg -1 yr -1 . All these estimates for the rate of dissolution of CaCO3 in the water column are summarized in Table 1. The results of Lerman and Lal [31] and of Li et al. [4] apply to the entire ocean and are reasonable when compared to the values for several discrete basins. It has already been shown that the laC and alkalinity data indicate that the western North Atlantic Ocean is anomalous in its high proportion of carbonate carbon compared to organic carbon. Hence it is to be expected that the dissolution rate calculated here is also slightly higher than previous oceanwide estimates.

5. Conclusions A particle of organic material entrained within the NADW mass undergoes oxidation at the rate of 0.10/amol of dissolved 02 kg -1 yr -1, producing 0.07/Jmol of CO2 kg -1 yr -1. An additional 0.12/Jmol of CO2 kg -1 yr -1 is produced by the dissolution of particles ofCaCO3. The observed decrease in dissolved 02 and increase in its 818 0 along the trajectory of the NADW indicates that the mechanism as well as the rate of utilization of the dissolved 02 are much the same as were observed earlier in the Pacific Ocean, and are attributed to bacterial respiration.

Acknowledgements I thank all the people in the GEOSECS Operations Group and the crew of the R/V "Knorr" for their assistance in collecting the samples. Ann-Marie Horowitz, Don Lingle and Dave Bos had the often frustrating job of keeping the vacuum extraction line working at sea. Rick Brill, Gale Sohl, and Alvin Sunn performed the laboratory analyses. The GEOSECS program at the University of Hawaii has been generously supported for a number of years by the International Decade of Ocean Exploration section of the National Science Foundation. This manuscript was written at the Physical Research Laboratory, Ahmedabad, India, thanks to a fellowship provided

by the Joint Indo-U.S. Subcommission of Cultural Affairs.

References P. Kroopnick, Respiration, photosynthesis, and oxygen isotope fractionation in oceanic surface water, Limnol. Oceanogr. 20 (1975) 988-992. P. Kroopnick and H. Craig, Oxygen isotope fractionation in dissolved oxygen in the deep sea, Earth Planet. Sci. Lett. 32 (1976) 375-388. H. Craig, Abyssal carbon and radiocarbon in the Pacific, J. Geophys. Res. 74 (1969) 5491-5506. Y. Li, T. Takahashi and W.S. Broecker, Degree of saturation of CaCO3 in the oceans, J. Geophys. Res. 74 (1969) 5507-5525. P. Kroopnick, The dissolved O2-CO2-1 aC system in the eastern equatorial Pacific, Deep-Sea Res. 21 (1974) 211227. K. Wyrtki, The oxygen minima in relation to oceanic circulation, Deep-Sea Res. 9 (1962) 11-23. H.H. Kuo and G. Veronis, Distribution of tracers in the deep oceans of the world, Deep-Sea Res. 17 (1970) 2946. M.E. Fiadeiro, Numerical modeling of tracer distributions in the deep Pacific Ocean, Ph.D. Dissertation, University of California at San Diego (1975). W.S. Broecker, T. Takahashi and Y.H. Li, Hydrography of the Central Atlantic, I. The two-degree discontinuity, Deep-Sea Res. 23 (1976) 1083-1104. 10 P. Kroopnick, The distribution of 13C in the Atlantic Ocean, Earth Planet. Sci. Lett. 49 (1980) 4 6 9 4 8 4 (this issue). 11 H.G. Ostlund, H.G. Dorsey and C.G. Rooth, GEOSEC8 North Atlantic radiocarbon and tritium results, Earth Planet. Sci. Lett. 23 (1974) 69-86. 12 M. Stuiver and S.W. Robinson, University of Washington GEOSECS North Atlantic carbon-14 results, Earth Planet. Sci. Lett. 23 (1974) 87-90. 13 M. Stuiver, The 14C distribution in West Atlantic abyssal waters, Earth Planet. Sci. Lett. 32 (1976) 322-330. 14 H. Craig, A scavengingmodel for trace elements in the deep sea, Earth Planet. Sci. Lett. 23 (1974) 149-159. 15 Y. Horibe, personal communication (1971). 16 Scripps Institution of Oceanography, Marine Technician's Handbook, Oxygen analysis, Sea Grant Publ. No. 9, Ref. No. 71-78 (1971). 17 P. Kroopnick and H. Craig, Atmospheric oxygen: isotopic composition and solubility fractionation, Science 175 (1972) 54-56. 18 D.E. Goodney and P. Kroopnick, Current techniques for the collection and processing of stable isotope data, in: Stable Isotopes in the Earth Sciences, B.W. Robinson, ed., DSIR Bull. 220 (Science Information Bulletin, Wellington, N.Z., 1973) 223-229.

496 19 H. Craig, Isotopic standards for carbon and oxygen and correction factors for mass-spectrometric analysis of carbon dioxide, Geochim. Cosmochim. Acta 12 (1957) 133-149. 20 P. Kroopnick, R.F. Weiss and H. Craig, Total CO2, 13C, and dissolved oxygen-18 at Geosecs II in the North Atlantic, Earth Planet. Sci. Lett. 16 (1972) 103-110. 21 G. Neumann and W.J. Pierson Jr., Principles of Physical Oceanography (Prentice-Hall, Englewood Cliffs, N.J., 1966). 22 GEOSECS Preliminary Reports, Legs 1 through 9, Atlantic Expeditions (GEOSECS Operations Group, Scripps Institution of Oceanography, La Jolla, Calif., 1972 1973). 23 T. Takahashi, P. Kaiteris, W.S. Broecker and A.E. Bainbridge, An evaluation of the apparent dissociation constants of carbonic acid in sea water, Earth Planet. Sci. Lett. 32 (1976) 458-467. 24 H. Craig, Abyssal carbon-13 in the South Pacific, J. Geophys. Res. 75 (1970) 691-695. 25 A.C. Redfield, B.H. Ketchum and F.A. Richards, The influence of organisms on the composition of seawater, in: The Sea, M.N. Hill, ed. (Interscience, New York, N.Y., 1963).

26 C. Culberson and R.M. Pytkowicz, 02-total CO2 correlation in the eastern Pacific Ocean, J. Oceanogr. Soc. Jpn. 26 (1970) 25-30. 27 P.M. Williams, J.A. McGowan and M. Stuiver, Bomb carbon-14 in deep-sea organisms, Nature 227 (1970) 3 7 5 376. 28 B.L.K. Somayajulu, D. Lal and S. Kusumgar, Man-made carbon-14 in deep Pacific waters: transport by biological skeletal material, Science 166 (1966) 1397-1399. 29 P.M. Williams, H. Oescher and P. Kinny, Natural radiocarbon activity of the dissolved organic carbon in the northeastern Pacific Ocean, Nature 224 (1969) 256-258. 30 S. Krishnaswami, D. Lal and B.L.K. Somayajulu, Investigations of gram quantities of Atlantic and Pacific surface particles, Earth Planet. Sci. Lett. 32 (1976) 403 419. 31 A. Lerman and D. Lal, Regeneration rates in the ocean, Am. J. Sci. 227 (1977) 238 258. 32 A. Lerman and D. Lal, personal communication (1978). 33 G.A. Riley, Oxygen, phosphate, and nitrate in the Atlantic Ocean, Bull. Bingham Oeeanogr. Coll. 13-14 (1951 1954) 1-126.

DEPTH (m) 28 104 182 274 464 575 813 1083 1376 1575 1725 2356 2356 2595 2696 3089 3267 3368 3630 3896

OXY. (pm/kg) 273 245 220 250 222 242 276 271 277 277 277 277 278 278 278 280 282 283 286 289

O-18 (o/oo) -0.i 2.4 1.6 1.6 1.4 1.5 3.9 1.6 -0.2 2.4 0.3 -1.5 1.3 i.i 3.0 2.5 2.7 0.9 -I.i 2.6

DEPTH (m) 58 159 361 1291 2160 2363 2639 2729 2929 3064

OXY. (~m/kg) 309 292 292 279 277 277 277 277 279 284

O-18 (o/oo) 0.6 1.3 1.2 2.4 1.6 2.3 2.0 2.5 2.2 2.4

401 403 405 407 409 411 41]

ID

DEPTH (m) 7 57 118 209 308 411 505

OXY. O-18 (~m/kg) (o/oo) 308 -0.3 289 i.i 272 1.9 279 1.8 283 1.6 287 1.6 286 17

Sta. ii 63.50 N 35.18 W Depth 2406 8, 5,1972

503 505 508 518 104 106 108 110 112 116

ID

Sta. 5 56.90 N 42.56 W Depth 3391 8, 1,1972

302 304 306 308 312 314 316 318 320 702 322 708 707 710 711 715 716 718 720 722

ID

Sta. 3 51.00 N 42.98 W Depth 4293 7,29,1972

6 x80 and selected hydrographic data for the Atlantic GEOSECS expedition

Appendix 608 709 810 861 1005 1105 1307 1506 1757 1862 2152 2350

288 286 289 287 287 282 278 276 277 279 289 303

1.8 1.4 1.8 2.0 1.9 2.0 2.4 2.6 2.5 2.5 2.6 1.6

DEPTH (m) 1 75 303 549 1297 1601 2502 2712 3102 3585

OXY. (pm/kg) 320 348 326 315 314 317 321 321 321 321

O-18 (o/oo) 0.2 0.6 1.9 1.9 2.0 2.1 1.9 0.8 1.7 1.7

DEPTH (m) 3 77 241 401 560 756 1147 2174 2523 3413

OXY. O-18 (~m/kg) (o/oo) 280 -0.8 270 1.9 305 1.5 307 0.6 308 0.2 305 2.5 301 2.9 303 2.5 303 2.8 301 3.0

701 703 709 715 719 401 721

ID

DEPTH (m) 1 70 401 674 888 980 1048

OXY. (~m/kg) 274 269 261 239 224 236 242

O-18 (o/oo) 3.7 i.i 1.5 2.8 3.8 3.7 3.7

Sta. 23 60.38 N 18.66 W D e p t h 2513 8,29,1972

602 604 608 611 616 619 303 311 317 323

ID

Sta. 19 64.17 N 5.59 W D e p t h 3495 8,25,1972

601 603 608 612 621 403 411 415 419 423

ID

Sta. 17 74.90 N 1.20 W Depth 3701 8,19,1972

415 417 419 421 102 104 106 108 ii0 112 116 120 1120 1422 1584 1939 2094 2310 2482

251 272 275 277 275 275 277

3.2 2.5 2.5 2.3 2.6 2.6 2.7

DEPTH (m) 16 24 63 89 1910 2086 2265 2461

OXY. O-18 (pm/kg) (o/oo) 292 2.9 291 3.1 283 2.5 283 2.8 276 -3.2 276 1.6 276 2.0 276 2.5

DEPTH (m) 4 18 29 47 68 98 2676 3580

OXY. O-18 (~m/kg) (o/oo) 220 0.I 221 -0.I 247 -0.2 232 0.9 210 1.7 207 2.9 272 i.i 276 -1.8

1001 i011 1015 1018 801 811 816 820 522 824

ID

DEPTH (m) 5 591 747 899 1492 2995 3589 4191 4662 4685

OXY. O-18 (~m/kg) (o/oo) 202 0.8 134 7.3 130 --0.2 143 7.1 239 3.6 259 4.9 260 3.3 258 3.3 254 3.2 254 3.2

Sta. 34 18.03 N 53.95 W D e p t h 4735 9,29,•972

315 318 319 320 321 322 119 Iii

ID

Sta. 26 45.01 N 42.09 W D e p t h 4708 9,11,1972

616 617 618 619 321 322 323 324

ID

Sta. 24 53.83 N 33.56 W D e p t h 2524 9, 8,1972

403 407 410 416 418 421 423

4~ ~D

DEPTH (m) 15 39 115 227 278 651 849 1299 2840 3239 3601 4200 4751 5018

OXY. (~m/kg) 199 214 168 105 117 126 136 205 256 254 264 265 250 245

O-18 (o/oo) -3.9 -5.8 -1.2 7.4 5.2 7.2 5.8 3.1 I.i 1.8 1.9 0.7 -0.5 2.5

DEPTH (m) 30 82 153 274 594 759 1090 1357 1576 2151 2750 3350 3898 4346 4791 5069

OXY. O-18 (~m/kg) (o/oo) 207 -0.8 203 -0.2 137 5.0 117 7.7 143 7.1 163 5.4 172 5.5 207 4.5 237 3.0 257 2.9 255 2.2 261 2.4 259 1.3 234 3.8 227 2.5 225 4.7

901 905 909 915 919 602 604 606

ID

DEPTH (m) 13 158 323 527 805 1336 1632 1934

OXY. O-18 (~Im/kg) (o/oo) 220 -2.9 211 -0.7 217 1.6 207 3.9 220 1.6 192 1.7 218 5.5 248 -2.8

Sta. 58 27.00 S 37.01 W Depth 4602 11,18,1972

902 904 906 908 912 916 920 923 102 105 108 iii 116 119 122 124

ID

sta. 48 3.98 S 29.06 W Depth 5075 10,28,1972

601 602 604 606 607 611 615 619 308 310 315 319 322 324

ID

Sta. 37 12.02 N 51.03 W Depth 5036 10,14,1972

Appendix (continued) 2235 2912 3943 4248 4578

253 256 230 225 225

1.3 -0.i 1.5 1.7 2.1

DEPTH (m) 207 662 1504 2651 4472 5193

OXY. (~m/kg) 270 255 184 235 225 226

O-18 (o/oo) 0.8 2.7 5.2 3.1 4.8 2.6

DEPTH (m) i0 151 251 405 559 658 1438 1741 2758 2982 3444 4424 4839 5804

OXY. (pm/kg) 298 274 263 274 264 254 184 183 226 217 209 218 224 227

0-18 (o/oo) -3.2 -1.9 -0.0 0.8 -1.8 2.7 3.3 6.2 2.1 2.8 2.7 3.8 2.8 2.7

DEPTH (m) ii 92 391 1657 3303 3420

OXY. (pm/kg) 325 320 204 188 218 220

O-18 (o/oo) -0.2 -0.5 4.4 5.3 3.9 4.3

501 505

ID

DEPTH (m) 8 126

OXY. (~m/kg) 335 330

O-18 (o/oo) -0.6 !.0

Sta. 74 55.01 S 50.04 W Depth 4144 12,18,1972

401 404 411 103 116 122

ID

Sta. 69 52.58 S 46.37 W Depth 3434 12,16,1972

401 405 407 410 415 416 423 102 107 108 ii0 118 120 124

ID

Sta. 67 44.98 S 51.01W D e p t h 5820 12,10,1972

906 912 923 708 719 723

ID

Sta. 64 39.18 S 48.59 w Depth 5357 12, 7,1972

608 612 620 622 624 2358 3346 3991

308 461 603 731 1205 1990 202 213 222

233 200 180 176 179 194 5.8 5.1 4.3

3.7 3.2 6.5 6.2 6.6 6.5

DEPTH (m) 96 500 700 1223 2116 2714

OXY. 0-18 (~m/kg) (o/oo) 336 2.1 178 0.i 179 -0.3 190 -0.4 203 0.2 210 0.i

DEPTH (m) 1 40 95 336 648 1299 2012 2402 2812 3213 3413 3862 5305

OXY. O-18 (pm/kg) (o/oo) 348 1.2 345 -0.i 315 1.2 186 7.4 195 -1.0 215 5.1 229 4.5 235 1.2 239 1.8 243 -4.8 245 2.3 249 -0.6 257 3.4

409 410 411 617 621 418 420 103 107 119

ID

DEPTH (m) Ii 49 144 384 745 1197 1446 2292 2891 3796

OXY. O-18 (pm/kg) (o/oo) 266 0.9 237 1.5 237 -1.6 233 7.7 222 3.6 179 5.1 176 5.8 218 4.0 226 4.2 218 4.9

Sta. 93 41.82 S .18.39 E D e p t h 4942 2, 2,1973

401 403 406 410 415 420 212 215 216 217 218 220 224

ID

Sta. 89 59.95 S 0.07 W D e p t h 5359 1,24,1973

505 109 517 522 115 119

ID

Sta. 78 60.98 S 63.00 W D e p t h 3706 i, 3,1973

112 1!9 123

509 515 519 521 105 109 DEPTH (m) 5 29 132 224 364 453 612 1054 1559 2061 2547 3342 4439 4563

OXY. (~m/kg) 217 221 228 191 131 iii ii0 171 206 231 238 235 220 218

O-18 (o/oo) 0.5 0.0 1.9 5.0 5.1 6.8 7.4 6.2 4.5 3.7 3.7 3.8 3.5 4.9

507 511 523 205 211

ID

DEPTH (m) 103 315 1609 2784 4439

OXY. (pm/kg) 109 97 240 246 249

O-18 (o/oo) 9.8 8.9 0.8 2.8 2.8

Sta.lll 2.00 N 13.99 W D e p t h 5153 3, 1,1973

601 603 607 609 611 612 615 619 621 623 301 305 310 317

ID

Sta.103 24.00 S 8.46 E D e p t h 4660 2,17,1973