The carbon dioxide system in the northwestern Indian Ocean during south–west monsoon

Marine Chemistry 64 Ž1999. 315–336

The carbon dioxide system in the northwestern Indian Ocean during south–west monsoon Ludger Mintrop a

a,b,)

a , Arne Kortzinger , Jan C. Duinker ¨

a

Institute of Marine Research, Department of Marine Chemistry, Dusternbrooker Weg 20, 24105 Kiel, Germany ¨ b UniÕersity of Bremen, Department of Geosciences, P.O. Box 330440, 28334 Bremen, Germany Received 9 October 1997; accepted 12 October 1998

Abstract Data on the carbonate system of the Northwestern Indian Ocean obtained on a cruise of F.S. Meteor during SW monsoon in JulyrAugust 1995 were compared with those of George et al. wGeorge, M.D., Kumar, M.D., Naqvi, S.W.A., Banerjee, S., Narvekar, P.V., de Sousa, S.N., Jayakumar, D.A., 1994. A study of the carbon dioxide system in the northern Indian Ocean during premonsoon. Mar. Chem. 47, 243–254x collected during intermonsoon. In general, deep water values agreed well between the two expeditions. Surface waters, however, showed a substantial increase in dissolved inorganic carbon Ž C T . in the coastal regions due to strong upwelling in the SW monsoon. This was also accompanied by very high CO 2 partial pressures in surface waters. The north–south gradients in vertical profiles of the measured parameters in the Arabian Sea are discussed by comparing profiles from the oligotrophic equatorial region with those from the highly productive central Arabian Sea. The effect of denitrification on regenerated C T and A T is minor, with contributions of - 9 and - 8 mmol kgy1, respectively, to the total amount regenerated also utilizing oxygen. The dissolution of biogenic carbonates is discussed; different approaches to define the depth, where the dissolution starts ŽlysoclineŽs., carbonate critical depth ŽCCrD.., are compared together with the calculation of saturation depth from carbonate concentrations. It is shown, that small differences in measured C T and A T Žfound between our data and those measured during GEOSECS. and different calculation approaches to the CO 2 system Ždifferent dissociation constants for species involved and taking into account phosphate and silicate concentrations. can produce pronounced differences in the calculated saturation depths. However, C T and A T data suggest substantial dissolution of biogenic carbonate in the water column even above the calcite lysocline, irrespective of the procedures followed to calculate this horizon. q 1999 Elsevier Science B.V. All rights reserved. Keywords: Indian Ocean; monsoon; carbon dioxide; calcite solution; denitrification

1. Introduction The Indian Ocean, especially the Arabian Sea, is an oceanic region influenced by a strong seasonal )

Corresponding author. Institute of Marine Research, Department of Marine Chemistry, Dusternbrooker Weg 20, 24105 Kiel, ¨ Germany. Tel.: q49-431-597-4023; Fax: q49-431-565-876; E-mail: [email protected]

cycle resulting from the monsoonal system. Also, areas of very high productivity are located here. A further prominent feature of the northern Arabian Sea is a zone of very low oxygen content at intermediate depth. These exceptional characteristics have implications to the cycling of carbon that invited several investigations on the carbonate system in the last 30 years Že.g., Sen Gupta and Pylee, 1968; Sen Gupta et al., 1976; Kumar et al., 1992; Anderson and

0304-4203r99r$ - see front matter q 1999 Elsevier Science B.V. All rights reserved. PII: S 0 3 0 4 - 4 2 0 3 Ž 9 8 . 0 0 0 8 9 - 9

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Dyrssen, 1994; Naqvi et al., 1996.. In 1994, George et al. presented a study of the carbonate system during premonsoon against which we compare our data obtained recently on a cruise during SW monsoon. We will discuss the effects on the carbonate system resulting from the seasonal cycle, from the Žmeridional. gradient in productivity in the Arabian Sea and from the formal incorporation of denitrification in equations describing carbonrnutrientroxygen relations. Another aspect addressed in this paper is the dissolution of biogenic carbonate. The role of the ocean in the uptake of CO 2 that is released into the atmosphere by burning of fossil fuels is a subject of many investigations in recent years. The interaction of the anthropogenic CO 2 with the sediment will be the final step in the equilibration between the atmosphere and the ocean. That eventually reduces the atmospheric pCO 2 substantially. The lysocline depth from where the carbonate minerals Žcalcite, aragonite. start dissolving, differs between the ocean basins, being the deepest in the North Atlantic and the shallowest in the Pacific Ocean. In both these oceanic regions the anthropogenic CO 2 may have already reached the respective lysoclines. Though the penetration of anthropogenic CO 2 into the North Pacific is not very deep, the lysocline is very shallow ŽChen et al., 1988.. In the North Atlantic despite the deep lysocline, anthropogenic CO 2 has already reached the bottom of the Žwestern. basins ŽChen, 1982; Gruber et al., 1996; Kortzinger et al., 1998.. ¨ For the Indian Ocean, a situation somewhere in between can be assumed. The dissolution of biogenic carbonates and the depth of the lysocline in the Indian Ocean shall be addressed here using the carbon data measured on this cruise. 2. Materials and methods Data were collected during cruise 32r5 of F.S. Meteor, leaving Mahe´ ŽSeychelles. on July 15 and arriving at Muscat ŽOman. on August 15, 1995. This cruise was a part of the German JGOFS ŽJoint Global Ocean Flux Study. contribution to the Arabian Sea Process Study. Stations were located on a transect initially heading north from the equator along 658E; at 158N the transect turned NW towards the Omani coast. The cruise track and the sampling

locations are shown in Fig. 1, including those of George et al. Ž1994. that we used for comparison. Water samples were collected with a CTD-rosette equipped with twenty-four 10 l-Niskin bottles. Total dissolved inorganic carbon Ž C T . and titration alkalinity Ž A T . were measured at 17 stations Ž430 samples. according to the standard operation procedure ŽSOP. outlined in DOE Ž1994.. Nutrients Žnitrite, nitrate, phosphate and silicate. and dissolved oxygen were measured in parallel. C T was measured by a coulometric method ŽSOMMA-analyzer, Johnson et al., 1993. that was calibrated using the certified reference material ŽCRM, batch 27, provided by A. Dickson ŽScripps Institution of Oceanography, La Jolla, CA, USA... A T was determined by potentiometric titration with hydrochloric acid. An open cell titration was carried out, using a calibrated pipette for delivery of the sample. The titration alkalinity was calculated from the titration curve Žusually 27 points. by adopting a curve fitting procedure ŽCampbell and Millero, 1994.. The system was calibrated with the CRMs, using the alkalinity of 2214.2 mmol kgy1 , measured for batch 27 by D. Campbell Ž1995, personal communication.; a recent determination from archived samples of this batch gave 2214.89 mmol kgy1 ŽA. Dickson, 1996, personal communication.. Precisions of C T and A T were determined by measuring duplicate samples at least once for every cast and by comparing the values measured for the CRMs throughout the cruise. Precisions for duplicate samples were "0.5 and "1.0 mmol kgy1 for C T and A T , respectively, while values for the CRMs agreed within "1.5 and "2.5 mmol kgy1 , respectively. At the beginning of the cruise a batch of surface seawater was collected and poisoned with mercuric chloride. This was used as a secondary standard for frequent measurements during the cruise to monitor the precision of the alkalinity titration. Analysis of this secondary standard yielded the best precision of "0.5 mmol kgy1 . Samples of the acid prepared volumetrically during the cruise were taken to the laboratory for recalibration after the cruise. The pCO 2 of surface water was measured continuously along the track by an automated device based on infrared detection ŽKortzinger et al., 1996.. The ¨ equilibrator was fed with ; 2 l of seawater per minute by a submersible pump installed in the ‘moon

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Fig. 1. Map of the cruise track of F.S. Meteor, July–August 1995 with stations Žsquares.. Also shown are the stations Ždots. occupied by George et al. Ž1994. on board of FORV Sagar Sampada in February 1992.

pool’ of the ship. The pCO 2 in air was measured at preset time intervals with the same device. For this purpose, air was pumped continuously from the ‘monkey’s island’ to the system. The system was calibrated with CO 2 standards in natural air at regular time intervals. These standards were calibrated by I. Levin, Institut fur ¨ Umweltphysik, Heidelberg, Germany, using primary standards from NOAA-CMDL, Boulder, CO, USA. Speciation of the carbonate system was calculated from C T and A T . A detailed discussion on the carbonate constants and computational methods used is given later Žsee Section 4.4.. The saturation levels of calcite and aragonite were evaluated from the calculated carbonate concentrations, the Ca concentration was derived from salinity Žsee also Section 4.4. and the stoichiometric solubility products were computed for in situ conditions. The solubility product was calculated using the equation ŽFeely et al., 1988.: X1 X1 log K sp s log K spŽ0. q Ž b 0 q b 1 P T q b 2rT . P S 0 .5

q c 0 P S q d 0 P S 1.5

Ž 1.

X1 where K sp is the stoichiometric solubility product at ambient pressure Ž p s 1 atm. as a function of salinity Ž S . and temperature ŽT, in Kelvin. and b 0 , b 1 , b 2 , c 0 and d 0 are different constants for calcite and X1 for aragonite. K spŽ0. is the value for pure water X1 Ž S s 0. and depends on temperature. To convert K sp Xp to K sp , the solubility product at in situ pressure Ž p, in atmospheres., the equation of Millero Ž1979. was used:

Xp X1 K sp s K sp P exp Ž y Ž DV P p q 0.5 P D K P p 2 . rRT .

Ž 2. where R is the gas constant, DV is the change in the molar volume and D K denotes the change in compressibility. Oxygen was measured by Winkler titration and nutrients were determined using standard photometric procedures with an autoanalyzer Žconstructed at Institut fur ¨ Meereskunde Kiel, Germany.. Apparent oxygen utilization ŽAOU. was calculated using the

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Fig. 2. Isoplots of Ža. total dissolved inorganic carbon Ž C T ., Žb. total alkalinity Ž A T . and Žc. salinity normalized total alkalinity Žn A T . along the transect from the equator to the Omani coast. Values are in micromoles per kilogram Žmmol kgy1 ..

equations for oxygen saturation given by Weiss Ž1970..

3. Results 3.1. Total dissolÕed inorganic carbon and alkalinity Isoplots of C T and A T along the transect are presented in Fig. 2a and b, respectively. A northward increase of C T is observed at the surface and at all depth levels. Surface values vary, but generally increase from below 2000 mmol kgy1 to around 2100 mmol kgy1 . Substantially higher values are associated with upwelling near the Omani coast. Deep water values, below 2000 m, increased from 2300 mmol kgy1 to 2360 mmol kgy1 along the transect. A similar pattern is found for alkalinity with low values of ; 2300 mmol kgy1 at the surface at the

southern end of the transect and of ; 2350 mmol kgy1 at the northern end. Maximum values Žup to 2380 mmol kgy1 . were measured between 12 and 148N. Deep water values increase from south Ž2400 mmol kgy1 . to north Ž2460 mmol kgy1 .. Alkalinity in Fig. 2b merely reflects the salinity distribution since its normalization to S s 35 levels out the surface maximum leading to a smooth and gradual increase in normalized alkalinity Žn A T . from - 2290 mmol kgy1 at the surface to 2350 mmol kgy1 at 1000 m in the entire study area ŽFig. 2c.. 3.2. Oxygen, AOU, nitrate reduction and resulting pCO2 A characteristic feature of the Arabian Sea is the existence of a broad depth horizon with very low oxygen concentrations Žoxygen minimum zone, OMZ., extending from below the thermocline to

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approximately 1000–1500 m. The lack of oxygen in this zone leads to the consumption of nitrate for the oxidation of organic carbon in sinking particles. Denitrification is thus a common feature in this oceanic region ŽDeuser et al., 1978; Sen Gupta and Naqvi, 1984.. The low oxygen zone Ž- 20 mmol kgy1 . is evident in the region north of 128N in the 200–1200 m range ŽFig. 3a.. This feature is also reflected in high values of the calculated AOU ŽFig. 3b.. The nitrate consumed during the oxidation of organic matter can be estimated by using the relation between measured nitrate and phosphate. We introduce here the term ANU, which is defined as: ANU s r N :P P PO43y

y NOy 3

meas .

meas .

Ž 3.

involving the measured phosphate and nitrate concentrations and the elemental ratio of nitrogen to phosphorus, r N:P . The theoretical N:P ratio would be 16 according to Redfield et al. Ž1963.. However, we used the ratio of 14.9 here, that was found as a mean by using all our data below the thermocline, but outside the OMZ. Eq. Ž3. indicates a rather simple qualitative description of the process but allows to identify the main denitrification areas Žnorth of 148N at 200–600 m, Fig. 3c.. The calculated nitrate deficit amounted to 8 mmol kgy1 , comparing favorably with the values of ) 6 mmol kgy1 found by Deuser et al. Ž1978. and 9–11 mmol kgy1 by Naqvi and Sen Gupta Ž1985. for their station 2690 Žclose to our station 414.. The depth of maximum nitrate deficit lies well above that of maximum in AOU Žsee Fig. 3b. in the upper part of the OMZ. Denitrification can be quantified by other methods ŽDeuser et al., 1978; Naqvi et al., 1982; Naqvi and Sen Gupta, 1985. as well. These, however, need analytical methods optimized for low level determinations Žoxygen, nitrate., additional measurements Žnitrite. and also imply a number of assumptions Že.g., nitratersalinity relations, Redfield ratios for O–C–N–P, calculation of ‘NO’, etc.., some of those being discussed controversially in literature for the area of investigation. However, we think that Eq. Ž3.

is sufficient to determine the order of magnitude of this process Žusing measured instead of preformed concentrations implies the assumption that denitrification only starts after oxic decomposition of organic matter had increased the nutrient concentrations according to ‘Redfield’ stoichiometry, see Section 4.2.. The remineralization of organic matter by oxygen and nitrate leads to very high CO 2 partial pressures Žderived from C T and A T . in the OMZ. Peak values of more than 1100 matm were calculated. These high values prevail at rather shallow depths ŽFig. 3d.. In surface waters, the continuously measured pCO 2 increased steadily along the cruise track from 370–380 matm Žatmospheric level: 350 matm. in the oligotrophic waters Žthe first—northbound—part of the transect. to values well above 400 matm in the eutrophic waters, i.e., after the track headed northwest. A steep increase to more than 700 matm was observed in the Omani coastal area. These high surface pCO 2 values were a result of coastal upwelling of subsurface waters. Fig. 4 shows a comparison of C T profiles in the upper 400 m of the water column with the surface pCO 2 profiles measured through the underway system, while approaching the Omani coast along the cruise track. We found similar surface C T values as in the upwelling areas Žwith pCO 2 ) 600 matm. at a depth of about 120–150 m outside this region otherwise. For example, the same pCO 2 Ž640 matm. as found at the surface at station 438 close to the coast is noticed at 130 m in the profile of calculated pCO 2 at station 430. Other water mass properties like temperature and salinity support these findings. Further results from the pCO 2 measurements on this cruise have been published in the work of Kortzinger et al. Ž1997.. ¨ 3.3. Carbonate dissolution The dissolution of calcite and aragonite depends on their solubility products, that change with temperature and pressure and on the in situ concentrations of carbonate and calcium ions. Since calcium concentrations are only slightly variable in the oceans,

Fig. 3. Same as in Fig. 2 but for: Ža. oxygen concentration, Žb. AOU, Žc. ‘apparent nitrate utilization’ ŽANU, see text. and Žd. calculated partial pressure of CO 2 at in situ conditions Ž pCO 2 .. Values are in micromoles per kilogram Žmmol kgy1 . for Ža. to Žc. and in microatmospheres Žmatm. for Žd..

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L. Mintrop et al.r Marine Chemistry 64 (1999) 315–336

Fig. 3 Žcontinued..

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Fig. 4. Comparison between C T Žmmol kgy1 . in the upper water column and surface pCO 2 Žmatm. along a part of the transect approaching the Omani coast.

the carbonate concentration determines the extent of dissolution of biogenic carbonate. We calculated carbonate concentrations at in situ conditions from the measured A T and C T ; the calcium concentrations were derived from salinity, assuming conservative behavior of Ca Ždetails on the calculation and the errors involved will be given in Section 4.4.. Isoplots of the degree of saturation for calcite Ž V C . and aragonite Ž VA . are shown in Fig. 5a and b, respectively. The saturation horizon Ž VA s 100%. of aragonite was found at about 600 " 100 m along the transect. For the samples deeper than 4500 m saturation of aragonite was as low as 50%. The

calcite saturation horizon is found at about 3000– 3200 m, with little variation along the track. The saturation horizons found are deeper than those reported by other authors. For instance, Naqvi and Reddy Ž1979. calculated the saturation horizon for calcite to shoal from 2600 m at the equator to about 1000 m in the northern Arabian Sea. Naqvi and Naik Ž1983. found a crossover from calcite supersaturation to undersaturation between 2500 and 3000 m. George et al. Ž1994. calculated the saturation horizon for aragonite at 500 m Žso only slightly shallower than our results. and for calcite at 1500 m for the Arabian Sea. Extremely shallow values are

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Fig. 5. Same as in Fig. 2 but for calculated degree of saturation Žin percent Ž%.. for Ža. calcite and Žb. aragonite.

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found by Kumar et al. Ž1992. with 200 m for the aragonite and 600 m for the calcite saturation horizons, respectively. The discrepancy between our values and reported values will also be discussed in Section 4.4.

4. Discussion 4.1. Distribution of total dissolÕed inorganic carbon and alkalinity George et al. Ž1994. presented C T and pH data for a transect running parallel to ours, but shifted NE by about 200 nautical miles Žn.m... They compared two stations, a southern one Ža2510. located 150 n.m. east of our station 404 and a northern station Ža2499. at 208N, 658E, about 250 n.m. NE of our station 423 Žsee Fig. 1 for locations.. They showed isoplots of C T along their transect and calculated pCO 2 and carbonate concentrations from their C T and pH measurements. Their data were collected in February, during the premonsoon period. Since the seasonal influence of the monsoon is unlikely to penetrate far below the thermocline, similar values can be expected at depths below several hundred meters. In general, the results presented by George et al. are very similar to ours. The same trends with latitude and depth were observed for C T and calculated pCO 2 . While their southern station should be comparable to our station at similar latitude, their northern station could be expected to show higher values resulting from increased remineralization within the water column. This may be concluded from the data of Naqvi and Sen Gupta Ž1985. who showed increasing values of the depth integrated nitrate deficit along their NW-transect nearly passing the two stations under consideration. Unfortunately, a direct comparison of the measured values is not possible since George et al. gave their concentrations in micromoles per liter Žmmol ly1 . without further information about the density of their samples. We will restrict the comparison to the observed gradients therefore. A northward increase in deep water C T of 30 mmol kgy1 from our stations 404 to 460 was found, while George et al. observed a steeper gradient between their stations 2510 and 2499 of 53 mmol

325

ly1 . The latter station is located further north than our stations, thus both datasets support a N–S gradient of deep water C T in the Arabian Sea. A seasonal effect is expected for surface values. George et al. found a moderate northward increase of C T at the surface by 56 mmol ly1 . Our surface values initially increased from the equator towards north from 1941 to 1985 mmol kgy1 at station 400 and then increased further to values around 2040 mmol kgy1 at the stations 414 to 430. So far, both data sets agree fairly well, when the increase from stations 400 to 430 Ž55 mmol kgy1 . is considered. The seasonal effect was conspicuous through a steep increase towards peak values of up to 2150 mmol kgy1 approaching the Omani coast, accompanied with high surface pCO 2 values Žup to 750 matm. caused by monsoonal upwelling. Two of our stations were located close to locations that had been sampled during the GEOSECSExpedition in 1979: one station was located near the Carlsberg Ridge Žour station a395, hereafter referred to as station M 395 and the GEOSECS station a418, hereafter G 418, at 6800 X Nr65800 X E and 6811X Nr64825X E, respectively.. The second station pair is situated further north in the Arabian Basin ŽM 411 and G 417, at 13800 X Nr65800 X E and 12858X Nr64828X E, respectively.. We compared C T and A T values of both station pairs. We excluded values shallower than 500 m to avoid the influence of seasonal variation and compared samples taken at nearly similar depth Žsame pressure within "2%.. For depths not sampled on METEOR, values were linearly interpolated from neighboring samples and the mean difference was calculated for the two sets of stations. For the pair M 411rG 417, A T values are slightly lower at G 417 by 3.5 " 4.7 mmol kgy1 while C T is higher by 16.8 " 6.4 mmol kgy1 . The latter is reduced to 1.8 mmol kgy1 after subtraction of 15 mmol kgy1 from the G 417 C T values, as recommended by Takahashi Ž1983. for the Indian Ocean C T data. However, for the other pair of stations, M 395rG 418, the situation is different. A T on average is higher at G 418 than at M 395 by 5.7 " 2.5 mmol kgy1 , C T is higher by 25.9 " 5.2 mmol kgy1 and corrected C T Žafter subtraction of 15 mmol kgy1 . is still higher by 10.9 mmol kgy1 . So the differences for the second station pair ŽG 418, M 395. are higher by 9.2 mmol kgy1 and 9.1 mmol

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kgy1 for A T and C T , respectively, compared to the first station pair ŽG 417, M 411.. Both values Ž C T and A T . were determined from the same titration in GEOSECS while two independent methods were used for the two parameter on our cruise. So an offset Žof similar magnitude for both A T and C T . between the two GEOSECS stations, perhaps resulting from changing to a new acid, a new titration cell or from a new calibration of acid andror cell volume, might be an explanation. We will come back to this comparison later in Section 4.4. 4.2. Stoichiometric relations for the remineralization of carbon including denitrification The changes in oxygen andror nutrient concentrations and in alkalinity enable to estimate the contributions from particulate organic carbon Ž‘soft tissue pump’. and biogenic carbonate Ž‘carbonate pump’. to the oceanic carbon flux. This approach is a prerequisite in the computation of the anthropogenic CO 2-signal. However, the question arises whether or not processes like denitrification, which has to be taken into account for the Indian Ocean, will affect these relations significantly. The remineralization of organic carbon is often described using the P:N:C ratios found in marine plankton by Redfield et al. Ž1963. Ž‘Redfield-ratios’. by the equation:

Ž CH 2 O . 106 Ž NH 3 . 16 Ž H 3 PO4 . q138O2l106CO2 q 16HNO 3 q 122H 2 O q H 3 PO4

Ž 4.

leading to the appropriate changes in C T , A T , oxygen and nitrate concentrations:

Ž 5.

This formulation, however, does not include the process of denitrification. In an attempt to include this process theoretically, by assuming that the end products of this process will be carbon dioxide and nitrogen and by using the stoichiometry: 5 Ž CH 2 O . q 4HNO 3 l 5CO 2 q 2N2 q 7H 2 O

Ž 6.

and 5NH 3 q 3HNO 3 l 4N2 q 9H 2 O

Ž CH 2 O . 106 Ž NH 3 . 16 Ž H 3 PO4 . q 94.4HNO 3 l 106CO 2 q 55.2N2 q 177.2H 2 O q H 3 PO4

Ž 8. resulting in: DC T s 106; D A T s Ž 94.4 y 1 . s 93.4; DO 2 s 0; DNO 3 s y94.4.

Ž 7.

Ž 9.

The use of nitrate for the oxidation of organic matter will possibly start already before the oxygen is fully depleted, but, according to Naqvi et al. Ž1996., an O 2 concentration below 7 mmol kgy1 is necessary for the onset of denitrification. The combined effect of oxygen utilization and denitrification therefore can probably be described as result of consecutive processes. With AOU s yDO 2 , combination of Eqs. Ž4. and Ž8. leads to: DC T ŽPOC . s 106r138

AOU y 106r94.4

DNO 3

Ž 10 . with DC TŽP OC. representing the increase in C T solely due to the remineralization of particulate organic carbon. The change in alkalinity is composed of Ži. the contribution from the dissolution of biogenic carbonate: CaCO 3 l Ca2qq CO 32y

Ž 11 .

with D A TŽc arbonate. s 2 Žfor this reaction DC TŽcarbonate. s 1. and Žii. of the alkalinity change due to the reactions formulated in Eqs. Ž4. and Ž8. resulting in: D A T s D A T Žcarbonate. y 17r138

DC T s 106; D A T s y Ž 16 q 1 . s y17; DO 2 s y138; DNO 3 s 16

an equation similar to Eq. Ž4. can be formulated:

y 93.4r94.4

AOU

DNO 3

Ž 12 .

Using the term ANU Žsee Eq. Ž3., but rather using the ‘classical’ Redfield value of r N:P s 16 to remain consistent with the previous theoretical evaluation. to describe in a schematic way the effect of denitrification: ANU s 16 P w PO4 x meas .y w NO 3 x meas .s yDNO 3

Ž 13 . the change in C T due to both oxidation of organic tissue by oxygen and nitrate and by dissolution of

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biogenic carbonate can be summarized by the equation: DC T s 0.5 P D A T Žcarbonate. q 106r138 q 106r94.4

AOU

ANU

Ž 14 .

inserting Eq. Ž12. gives: DC T s 0.5 P D A T q 114.5r138 q 59.3r94.4

AOU

ANU

DC T s 0.5 P D A T q 0.83AOU q 0.63ANU.

Ž 15 . Ž 16 .

The first two terms on the right side of Eq. Ž16. are those introduced by Chen and Millero Ž1979. for calculating the effect of remineralization on C T . The third term is added here to include the denitrification processes. Eqs. Ž14. and Ž12. indicate that denitrification processes have a similar effect on C T Žincrease by 1.12 ANU. and on A T Žincrease by ; 1 ANU; the ANU-factor of 0.63 in Eq. Ž16. is smaller because the D A T-term is also affected by denitrification, see Eq. Ž12... Since samples with high ANU typically also have high AOU values, the overall effect is rather small ŽANU F 8 mmol kgy1 vs. AOU G 240 mmol kgy1 .. All these equations are based on the Redfield ratios and hence the factor 16 is used for the N:P ratio in Eq. Ž13.. However, the use of the ‘classical’ Redfield ratio for this kind of calculations has been put into question and modified ratios have been proposed Ždifferent ratios for distinct water masses, for certain ocean areas or also ratios varying with depth., mainly resulting in higher values for oxygen ŽBroecker et al., 1985; Takahashi et al., 1985; Anderson and Sarmiento, 1994.. In fact, only the C:N:P relations in Eq. Ž4. are based on Redfield’s measurements, the relation to oxygen results from a ‘theoretical’ composition of organic matter ŽŽCH 2 O.106ŽNH 3 .16 ŽH 3 PO4 .. and the moles of oxygen necessary for full oxidation to CO 2 according to stoichiometry. However, this formula implies a ‘compound’ of mainly carbohydrate ŽŽCH 2 O. n . type. Compounds with aliphatic chains Že.g., lipids. form a substantial part of organic tissue, involving a partially negatively charged C-atom, which thus requires the transfer of more than four electrons for the oxidation to CO 2 . The C:O 2 relation of 122 " 18:172 formulated by Broecker et al. Ž1985. for the Indian Ocean Žwhich is also close to the ratio of Anderson

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and Sarmiento Ž1994. of 117 " 14: 170 " 10. is nearly equivalent to 106:149, thus resulting in 11 additional moles of oxygen for the remineralization of the ‘theoretical organic tissue’. Calculating the remineralization of organic tissue from AOU results in values lower by 7.7% when the Broecker et al. factor 122:172 Žs 0.709. is used instead of 106:138 Ž0.768.. Since these relations are averages for whole oceans, it is not clear, whether they will describe the stoichiometry in specific areas like the Arabian Sea adequately. Sen Gupta et al. Ž1976. calculated a relation for –O 2 :C:N:P of 140:108:16:1 for the NW Indian Ocean, which is rather close to classical Redfield ratios. Kroopnick Ž1985. proposed a value of 0.75 Žs ˆ 106: 141. for the C:O 2 relation, that was found as an average for the world ocean and has widely been used since. Additionally, other uncertainties Že.g., degree of oxygen saturation deviating from 100% at the surface. will also contribute to the amount of carbon remineralized from organic tissue when calculated from AOU. Similar considerations as for the factor to multiply with AOU should affect also Ži. the definition of ANU Ždepending on the N:P ratio used. and Žii. the effect of denitrification on carbon remineralization, i.e., the factor of 0.63 in Eq. Ž16.. However, according to the equations derived above, the effect on DC T is only greater than the experimental error of the C T determination Ž"2 mmol kgy1 . for high ANU values. To conclude, for the calculation of the contribution of organic matter remineralization to the total dissolved inorganic carbon concentration, the effect of denitrification in most cases will be smaller than the errors associated within the correlation factor linking AOU and carbon remineralization. However, for a relative comparison of different stations that is done in the following, the deviations for different AOU-factors should not be relevant. 4.3. Organic Õs. inorganic remineralization in the Arabian sea, N–S gradient Due to the prevailing diverse regimes in the Arabian Sea, ranging from the oligotrophic zone in the equatorial region to the high productivity areas in the north, generally N–S gradients are observed in the parameters associated with biogeochemical pro-

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cesses. Two stations were selected to illustrate the reflection of this meridional gradient on the carbon system. The southern station M 395 is located near the Carlsberg Ridge at 68N and 658E, whereas the second, the northern station M 423, near the Omani coast in the Arabian Basin is at 178N and 628E. Station M 395 is representative of an oligotrophic regime Žsituated close to GEOSECS station G 418.. Station M 423 represents the area of strong oxygen depletion and denitrification. 4.3.1. Total dissolÕed inorganic carbon Vertical profiles of C T for the two stations ŽFig. 6a. show that at M 423 values at any depth are higher than at M 395. C T reaches a constant value below 1500 m at M 395 while a further increase towards the bottom is observed at M 423. To compare the dissolution of organic tissue at the two stations, DC TŽP OC. was calculated from the AOU term in Eq. Ž10.. The dotted line in Fig. 6a is the possible additional contribution from denitrification processes occurring at station M 423 resulting from the DNO 3-term in Eq. Ž10.. For station M 395 this contribution proved to be negligible and thus was omitted in the figure. DC TŽP OC. has a maximum between 600 and 1200 m at M 395, whereas a much broader maximum, between 200 and 1500 m occurs at M 423. Below 1000 m, DC TŽP OC. is higher at M 423 than at M 395 by about 15–20 mmol kgy1 ; on the other hand, the difference in C T between the two stations increases with depth, so processes other than organic tissue remineralization obviously are important. 4.3.2. Alkalinity Since salinity is different at both the stations in the upper 1000 m and alkalinity is known to be a conservative property mainly, the normalized values

Fig. 6. Ža. Comparison of depth profiles of C T Ždiamonds., contribution from soft tissue remineralization to C T Žcircles, the dotted line indicates the contribution from denitrification, see text. for stations 395 Žfilled symbols. and 423 Žopen symbols.. Žb. Depth profiles of n A T Ždiamonds. and contribution of carbonate dissolution calculated to a reference level of su s 27.70 Žapproximately 1500–1600 m. for both stations Žindicated by open and filled symbols as in Ža.; all concentrations are in micromoles per kilogram Žmmol kgy1 ...

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were compared to exclude the salinity effect ŽFig. 6b.. The normalized profiles are found to be nearly identical from the surface to ; 1500 m. Values for both profiles increased towards the bottom, but this increase is far more pronounced for station M 423. Explanation for the different deep water alkalinity could be: Ži. the prevalence of different water masses Žthus having different preformed Žnormalized. alkalinity. at the two stations; Žii. an enhanced dissolution of biogenic carbonate in deep and bottom waters at station M 423 as a result of enhanced supply from the surface layer Žhigher productivity. andror from the sediment Žmore corrosive.. The prominent deep water mass with high preformed alkalinity would be Antarctic Bottom Water ŽAABW.. Since this water enters the Indian Ocean from the south and penetrates north, a northward increase is unlikely. The high productivity in the north, compared to that at station M 395, favors the explanation Žii. which is also in accordance with increasing DC T against depth Žthat is only partly paralleled by AOU. between the two stations. A sediment source is less likely because this should be restricted to the deep boundary layer and would hardly explain the extension of the observed feature into the water column several hundred meters above the seafloor. Calculating directly the amount of dissolved biogenic carbonate from A T is difficult, since the preformed A T-values have to be taken into account. All water masses bear their individual preformed C T and A T signal. The water masses that form the water of the Arabian Sea below 1500 m have their origin from south of the equator, but several end members have to be considered. A detailed summary of the water masses in the Arabian Sea has been given by Shetye et al. Ž1994. and Kumar and Li Ž1996., illustrating the complexity of the water mass composition of the northern Indian Ocean. The origin and the circulation of the Deep and Bottom Waters in the Arabian Sea are still a matter of debate ŽWarren, 1992.. It is generally observed for most oceanic regions, that surface normalized alkalinity decreases with temperature. This apparent temperature dependence can be regarded as an artifact, caused, for instance for the southern hemisphere, by upwelling of alkalinity rich deep waters in the Žcold. Southern Oceans

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ŽGruber et al., 1996.. Therefore, in the Arabian Sea, alkalinity should increase with depth Ždecreasing temperature. due to the temperature dependence of preformed values which might mask the effect of carbonate dissolution on alkalinity. Moreover, the temperature correlation for preformed values may be different for water masses of different origins, as can be expected from the various surface data published in the literature for different oceanic regions. The temperature relations with salinity normalized alkalinity reported for the South Atlantic, Southern Ocean and South Indian Ocean mostly have a slope of around y4 to y5 mmol kgy1 8Cy1 ŽBroecker et al., 1985; Takahashi et al., 1985; Poisson and Chen, 1987; Chen, 1988; Chen, 1992.. In order to have a reasonable estimate of the dissolution of carbonate shells from A T we used the relation: DC T Žcarbonate. s 0.5 P Ž Dn A T q 17r138

AOU . .

Ž 17 . To avoid uncertainties resulting from strong temperature gradients, we defined a reference horizon at su s 27.70 Žwhich is at approximately 1600 and 1500 m depth at M 395 and M 423, respectively. where n A T is ; 2400 mmol kgy1 at both stations. The potential temperature Ž u . at this horizon is close to 48C along the transect and decreases by less than 2.48C from this level towards the bottom. For the temperature slope of preformed normalized alkalinity we assumed a value of y4.6 mmol kgy1 8Cy1 ŽPoisson and Chen Ž1987. for southern Indian Ocean.. Since AOU decreases below 1500 m, the dissolution of carbonate relative to the reference horizon was calculated according to: DC T Žcarbonate. s 0.5 P Ž n A T y 4.6 P Ž 4 y u . y 2400 q17r138

D AOU .

Ž 18 .

and is shown in Fig. 6b. The value calculated using Eq. Ž18. will only represent a part of the increased alkalinity below 1500 m that results from enhanced carbonate dissolution relative to that at the reference horizon and thus represent a lower limit, since carbonates likely also dissolve above 1500 m. However, the temperature gradient from 500 m, the shallowest depth found for the aragonite saturation horizon by George et al. Ž1994., to the bottom is more than 118C and the error introduced by the temperature

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adjustment would obscure the dissolution effect, if extrapolation is made to this larger depth range. At station M 395 carbonate is dissolved rather constant at about 16 mmol kgy1 below 2700 m, while at M 423 a steady increase is found up to 39 mmol kgy1 at 4000 m. So carbonate dissolution in the deep Arabian Sea reaches at least 12% Žfor M 395. to 24% Žfor M 423. of the values calculated for soft tissue remineralization, even if the dissolution above the reference horizon of 1500 m is ignored. For comparison, Chen and Wu Ž1993. reported a contribution of 25% to the C T increase from the dissolution of calcium carbonate near the Gulf of Aden. We also tested the approach of Gruber et al. Ž1996. to use salinity and ‘PO’ ŽBroecker, 1974. as independent variables to determine preformed alkalinity. However, we found only a poor correlation between PO and salinity normalized alkalinity Žn A T . in surface waters, while no relation could be established for preformed surface alkalinity with salinity and PO. This is not surprising, since the authors could not apply their approach successfully to ocean regions outside the North Atlantic. A more rigorous calculation of carbonate dissolution from alkalinity requires knowledge of preformed alkalinity values and their temperature dependence Žor their relation to another conservative parameter. for all water masses involved. This cannot be achieved with the limited present data set collected from the Arabian Sea and therefore is beyond the scope of this paper. 4.4. Carbonate dissolution As has been pointed out above, the dissolution of biogenic carbonate in the water column plays an important role in carbon cycling in the Arabian Sea even above the calcite lysocline. The dissolution depends on thermodynamic factors, that is the saturation state of the seawater and also on the dissolution kinetics. One or both aspects are considered in various formulations that are used to describe carbonate dissolution. The term ‘lysocline’ has been used differently in different approaches: Berger Ž1968. defined a foraminifera lysocline, that is the depth above which no sign of dissolution can be found microscopically on foraminifera tests. The Ž‘sedimentary’-.

lysocline was defined by Kolla et al. Ž1976. as the depth where a sharp increase in the percentage of calcite in surface sediments is observed. A kinetic approach is the definition of a ‘hydrographic’ lysocline as a depth where an abrupt change in the rate of dissolution of calcite in the water column occurs and is based on in situ experiments Ži.e., kinetic studies., which are scarce in the world ocean. Indirectly, the parameter D pH, the difference between the pH of the calcite–seawater equilibrium and the seawater pH, has been taken to define the lysocline depth Žthermodynamic approach.. Takahashi Ž1975. suggested a value of 0.08 as an indicator for a sharp increase in the dissolution rate. Berner and Wilde Ž1972. found D pH 0.08 to be equal to a value of V s 91%. We included the isoline of V s 91% derived from our data in Fig. 5a and b. The depth range for aragonite is 800–1000 m whereas for calcite it is approximately 3500–4000 m. The calculation of D pH from our data gave the value of 0.08 at slightly greater depth ranging from 3600 to ) 4500 m for calcite and from 900 to 1100 m for aragonite, respectively. Broecker and Takahashi Ž1978. introduced the concept of a critical carbonate concentration, based on a study on carbonate ion concentrations at lysocline depth in the Pacific and the Atlantic. Their data fits yielded the following equations: CO 32y

s 90 P exp Ž 0.16 P Ž D y 4 . . for calcite,

crit .

Ž 19 . CO 32y

crit .

s 120 P exp Ž 0.15 P Ž D y 4 . . for aragonite

Ž 20 .

with D being the lysocline depth in kilometers. Application of this concept gave a lysocline depth of about 3500 m for calcite and 500–600 m for aragonite, respectively, significantly shallower than the V s 91% and D pH s 0.08 approaches. The carbonate compensation depth is defined as the depth where the rate of carbonate supply to the sediment is equal to its rate of dissolution, i.e., no carbonate is preserved. The CCrD is systematically shallower and defined as the depth where the carbonate content of the sediment is - 10%. Kolla et al. Ž1976. summarized the calcium carbonate distribution in the sediments of the Indian Ocean and found the CCrD to occur at 4800 m in the Arabian Basin.

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The relationships between the saturation horizon Ži.e., where V s 100%., the lysoclineŽs. and the calcium carbonate compensation depth ŽCCD. have been pointed out by several workers ŽEdmond and Gieskes, 1970; Takahashi, 1975; Broecker and Takahashi, 1978; Chen et al., 1988.. The application of different concepts on our Indian Ocean data yield lysocline depths between 3400 and ) 4500 m for calcite and the range for aragonite was from 500 to 1700 m. Thus different approaches to describe the dissolution of carbonate yield different results for the North Indian Ocean, as would be expected from the different definitions, but none of the concepts would explain significant dissolution of biogenic calcite shells except for the bottom water region. As pointed out before Žsee Section 3.3. the saturation horizon Ž V s 100%. that we found and that is reported in the literature are different. In the following, we will focus on the errors associated with the calculation of carbonate saturation using the stoichiometric solubility product to find if any improvement is possible through a more rigorous approach. From the stoichiometric solubility product and the Ca-concentrations the equilibrium concentration of carbonate is calculated and V is derived by comparison with ‘measured’ carbonate concentration. The overall error of V therefore depends on the errors associated with Ži. the stoichiometric solubility product Žat in situ pressure, temperature and salinity., Žii. the estimated Ca concentration and Žiii. the ‘measured’ carbonate concentration. However, the latter is not measured but rather calculated from two of the measurable parameters C T , A T , pH, or pCO 2 . The error in evaluated carbonate content therefore is composed of the experimental errors in the measured variables and the uncertainties of the equilibrium constants used in the calculations. 4.4.1. The stoichiometric solubility product Different values and temperature relations are forX1 ŽMucci, 1983; Sass et al., 1983.. mulated for K spŽ0. Both values almost agree at 258C, but the deviation between these two formulations increase with decreasing temperature. Given the typical temperature profiles measured in the Arabian Sea, the difference in the calculated carbonate saturation concentration for calcite increases almost linearly from close to zero at 1000 m depth to more than 5 mmol kgy1

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Ž4.7%. at 4750 m. Since authors did not always state explicitly the equations they used, the choice of X1 K spŽ0. can explain differences in reported saturation concentrations. The overall accuracy of the stoichiometric solubility product depends also on the salinity and pressure relations. The equations result from laboratory experiments with rather few data points for the fit. Millero X1 Ž1979. estimated an error of up to 6.7% in K sp Žtemperature and salinity effect. and an uncertainty of 1 to 2.4% per km with respect to the pressure correction. 4.4.2. Calcium concentrations Calcium concentrations in seawater at open ocean salinities are about 10 mmol kgy1 . The calcium concentration is generally derived from salinity, since the CarS relation is rather constant Ž"1.5%. in the world ocean ŽCulkin and Cox, 1966.:

w Ca2q x s 2.938 P 10y4 P S

mol kgy1 .

Ž 21 .

However, in the Arabian Sea pronounced deviations from this relation have been found. Taking into account the value for the Ca Žgiven in g kgy1 .rchlorinity ratio of 0.02164 reported by Naqvi and Naik Ž1983., the relation between Ca and salinity becomes:

w Ca2q x s 2.989 P 10y4 P S

mol kgy1 .

Ž 22 .

For typical deep water salinities of 34.8, the difference is 178 mmol kgy1 ; however, this is only slightly above the range of Eq. Ž21. when an uncertainty of "1.5% Ži.e., "153 mmol kgy1 at S s 34.8. is assumed. The higher Ca concentration resulting from Eq. Ž22. would lower the calculated lysocline slightly. However, the uncertainty in the Ca concentration would only result in an error of less than "2%. 4.4.3. Carbonate concentration Two main factors affect the calculated carbonate concentration: the quality of the input parameters Ži.e., C T and A T . and the choice of constants Žand the consideration of dissolved nutrients. in the calculation. To check the influence of input parameters we used the C T and A T values from station M 411 compared to the values from G 417 Žwith C T from G 417 corrected by 15 mmol kgy1 , see Section 4.1. to calculate carbonate concentrations.

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The speciation of the CO 2-system is calculated from any two Žor more. of the measurable parameters C T , A T , pH and pCO 2 and the appropriate equilibrium constants. The latter have been experimentally determined by several investigators for various salinity and temperature ranges. Apart from the errors in the measured parameters themselves, the choice of the parameter pair and equilibrium constants shall also influence the results ŽStoll et al., 1993; Lee and Millero, 1995.. No standard procedure has been generally accepted yet. The determination of carbonate concentrations from A T and C T is done through the estimation of wHqx in situ ŽEdmond and Gieskes, 1970; Keir, 1979.. This is usually done using computer programs that allow for iterative adjustment of the results. The contributions of minor bases to alkalinity, at the previously calculated pH, is taken into account in the evaluation of carbonate alkalinity Ž A C .. Then the carbonate species can be derived from C T and A C using the dissociation constants for carbonic acid. We used the program of Lewis and Wallace Ž1998., hereafter abbreviated as LWP and these results were compared with the carbonate values calculated from GEOSECS data by Takahashi Ž1983., hereafter called the GP. To illustrate the differences in resulting carbonate concentrations, we recalculated the G 417 carbonate concentrations from A T and corrected C T values using LWP taking into account the nutrient concentrations given by Weiss et al. Ž1983.. The difference between both programs is that: GP calculates on the NBS-pH-scale using the carbonic acid constants of Mehrbach et al. Ž1973., the boric acid constant of Lyman ŽLi et al., 1969. and boron concentrations from Culkin Ž1965.. The effects of OH, phosphate and silicate concentrations on alkalinity and the non-ideal behavior of CO 2 are not accounted for in this program. LWP has different options for the choice of constants and pH-scales. We used the carbonic acid constants of Roy et al. Ž1993., the boric acid constant of Dickson Ž1990b., sulfate concentration estimated from Morris and Riley Ž1966., the sulfuric acid constant of Dickson Ž1990a.. Fluorine concentration is taken from Riley Ž1965., while the fluidic acid constant is from Dickson and Riley Ž1979.. LWP includes the contribution of hydroxide, phosphate and silicate to alkalinity Žconstants used

are from the work of Millero, 1995.. The total pH-scale was used. Fig. 7 shows the comparison of the vertical profiles of calculated carbonate concentrations. The carbonate values from LWP are generally lower Žby ca. 2 mmol kgy1 below 2000 m. than the data from GP ŽGP does not take into account nutrients, so ‘nn’ is added for ‘no nutrients’.. We repeated the calculations with LWP, this time setting the concentrations of phosphate and silicate to zero ŽLWP, nn.. Roughly, the carbonate alkalinity calculated by taking into account nutrients decreases by an amount equal to the phosphate concentration Žin micromoles per kilogram Žmmol kgy1 .. and by 0.01 times the silicate concentration. However, values from LWP come very close to the results from GP but are slightly higher, when the nutrients are ignored. This might indicate that the calculation is more dependent on

Fig. 7. Carbonate concentrations calculated for stations G 417 and M 411 by different program options Žsee text for details.. Also indicated are the saturation concentrations calculated after Mucci Ž1983. and Sass et al. Ž1983. for calcite and aragonite, respectively.

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inclusion of nutrients than the use of different acid constants. We then calculated the carbonate concentrations from the M 411 carbon and nutrient data using the same settings for LWP. Now, the carbonate concentrations are much higher than the G 417-LWP results and even higher than the results from G 417-GP. Not taking into account the nutrients for the M 411 data ŽLWP, nn. results in even higher carbonate concentrations. This example demonstrates that taking into account the nutrients has a substantial influence on calculated carbonate concentrations; either the choice of various dissociation constants for the species involved in the calculation plays a minor role or the different effects are levelled out favorably by coincidence for the example data chosen. However, even when C T and A T data for the same location differ only slightly, as was the case here, still substantially different carbonate ion concentrations Žhere 5–7 mmol kgy1 . may result from the calculation. At the moment we cannot decide which program produces the right value, but it is obvious, that the choice of the program gives considerably different results from the same set of carbon data and that even quite similar data from two different investigators can result in an uncertainty about the depth of the lysocline of several hundred meters. In regard of the larger scatter of the GEOSECS carbonate data Žsee Fig. 7., an improved situation can be expected to result from recent measurements under WOCE and JGOFS, which are likely to provide a better database due to refinement of analytical methods. Also included in Fig. 7 are the equilibrium carbonate concentrations calculated from the Ca-concentration Žusing Eq. Ž21.. and the stoichiometric solubility product, both using the formula given by Mucci Ž1983. and Sass et al. Ž1983., indicated by hatched and dotted lines, respectively. Further errors resulting from the formulation of the pressure dependence and a possible different CarS relation may even increase the uncertainty of the calculated equilibrium carbonate concentrations. To conclude, both the uncertainties in the thermodynamic constants and in the calculation of carbonate concentration from measured parameters add almost equally to the overall uncertainty, while the deviation from a linear Ca–salinity relation seems to

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play a minor role Žand could be accounted for by direct measurements of Ca-concentrations.. As a number of approaches are currently followed to link carbonation concentrations with carbonate concentrations in sediments Že.g., Archer, 1996. or to predict the buffer capacity of surface sediments for anthropogenic CO 2 increase by modeling, the attention of investigators should be drawn to the uncertainties discussed above. Apart from the thermodynamic equilibrium determined from vertical profiles of measured parameters, the dissolution of carbonates can be favored by conditions on a smaller scale, e.g., microenvironments where respiration of microorganisms can create locally high CO 2 levels around particles, that would possibly create a pH low enough to lead to calcite dissolution.

5. Conclusions Apart from the seasonally influenced surface layer, values for the CO 2 system agree well with those obtained by George et al. Ž1994. for their parallel track, that was slightly shifted NE compared to our transect. This shift can also explain their higher deep water values for C T at their northernmost station. The effect of the SW monsoon on surface C T became evident from the comparison. Very high C T values are found in the upper water column accompanied with upwelling, elevated up to 110 mmol kgy1 over intermonsoon levels. The most striking feature observed during monsoon is the extremely high pCO 2 measured near the Omani coast. The CO 2 characteristics of the upwelled water correspond to a source depth of approximately 120 m. From a theoretical point of view, the process of denitrification should only slightly influence the dissolved inorganic carbon concentration compared to the large contribution from the remineralization of organic matter by use of oxygen. Alkalinity presumably is increased by about the same amount Žin micromoles per kilogram Žmmol kgy1 .. as nitrate is reduced during the denitrification. Difficulties arise when the depth, below which thermodynamical dissolution of aragonite and calcite minerals occurs, has to be defined. Different concepts Žsaturation concentration, critical carbonate

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concentration, D pH, lysocline, CCD, CCrD. lead to different horizons. Obviously the calculation of carbonate concentration from measured CO 2 data requires more attention. Sources of errors are the quality of the C T and A T data, the calculation routines Žconstants used, considering nutrient concentrations. and the thermodynamic constants. Regarding the latter, new experiments to increase the range and the accuracy of the salinity, temperature and pressure dependencies of the stoichiometric solubility product seem to be highly desirable. Improvement of analytical techniques and the huge amount of high quality data collected in recent years under the JGOFS and WOCE programs certainly will improve the data situation significantly, as consistency checks ŽMillero et al., 1993; Clayton et al., 1995; Lee et al., 1996; Lee et al., 1997. of these data will help finding the appropriate choice of constants in the carbon system calculations Žwhere also nutrient concentrations should be accounted for.. Though the saturation horizon for calcite is found at considerable depth Žby any approach., dissolution of carbonate is obvious in the water column at least below approximately 1500 m. The reason for this remains to be identified: Ž1. The biogenic carbonate dissolved above the calcite lysocline is composed purely of aragonite; a significant contribution of this modification to biogenic carbonate has been proposed by Berger Ž1978.. Ž2. Carbonate dissolution within the sediment is responsible for the A T increase; at least for silicate the sediment has been identified as a source for the deep water ŽKumar and Li, 1996.. However, this effect is likely to be limited to near bottom layers. Ž3. Apart from thermodynamical equilibrium, calcite can be dissolved due to specific conditions in microenvironments associated with organic coating of the particles; it is not unlikely, that respiration of microorganisms can create high CO 2 levels Žand therefore low pH. within cellular structures, that would dissolve calcite ŽAnderson and Sarmiento, 1994.. Analyzing the contributions from the different water masses to the Arabian Sea CO 2-system will be essential to evaluate preformed values of total dissolved inorganic carbon and alkalinity. The recent international cruises in the Indian Ocean will provide a wealth of data for this purpose. So correcting A T

for preformed values will allow for a more rigorous calculation of carbonate dissolution.

Acknowledgements We thank captain and crew of F.S. Meteor for excellent cooperation during the cruise. Special thanks for perfect organization to chief scientist B. Zeitzschel. S. Schweinsberg and C. Reinecke skillfully performed the analysis on board. Critical revision by Dr. Kumar helped improve the manuscript substantially. The work was funded by the German Ministry of Research ŽBMBF. under grant no. 03F0160A and is part of the German JGOFS contribution.

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