waters and biological activities, the covariance of wind speed and ∆pCO2 should be taken into consideration in future studies.
To correct for interannual changes in the oceanic and atmospheric pCO2 values, the observed ∆pCO2 values have been normalized to a reference year 1990 as explained earlier. The flux values for high latitude oceans computed using the “full” atmospheric CO2 increase are about 10% greater in magnitude than the respective values obtained for the “half” effect ( Table 2 ), and hence the error due to this effect is about ±5%.
Errors due to the computational method has been estimated to be about 5 µatm for the mean global ∆pCO2. This corresponds to an error of up to 75% of the flux estimates made using a given gas transfer formulation.
∆pCO2 depends on the skin temperature of surface ocean water (49, 74), while the pCO2 has been evaluated at the bulk water temperature in this study. The effect of skin layer cooling on the global CO2 uptake has been estimated to be 0.1–0.6 Gt-C·yr−1 (75). However, the skin temperature may be higher or lower than the bulk water temperature depending upon meteorological and oceanic conditions, and the measurements are limited in space and time. Therefore, its effect on the global ∆pCO2 and flux has been neglected in this study.
A database for the sea–air pCO2 difference, ∆pCO2, has been assembled using about 250,000 observations made between 1960 and 1995 during 250 expeditions over the global oceans. Observations made in the equatorial Pacific during El Nino events have been excluded. In light of the sparseness of observations over large oceanic areas, the multiyear data have been corrected and combined to represent a single reference year of 1990. These observations have been organized into 4° latitude × 5° longitude × 1 day pixels for 365 days, and interpolated in space and time using a computational scheme based on the diffusive and advective transport of surface water (63).
On the basis of the global distribution of ∆pCO2 values thus computed, a global net ocean uptake of 0.60 to 1.34 Gt-C·yr−1 is obtained for 1990 using three different formulations for the gas transfer coefficient. This is similar to 1.6 ± 0.9 Gt-C·yr−1 estimated on the basis of 13C changes in the atmosphere and ocean (2), but is smaller than about 2 Gt-C·yr−1 based on various ocean–atmosphere perturbation models (3 , 4, 5, 6, 7 and 8). However, it is greater than the estimates based on the atmospheric CO2 distribution and mass balance (1).
The Pacific equatorial belt is the largest oceanic CO2 source to the atmosphere. The temperate oceanic areas of the both hemispheres are the most important sinks, and their uptake fluxes exceed those of high latitude oceans (poleward of the 50° parallel) by a factor of 2 to 3. Among the four ocean basins, the Atlantic Ocean (north of 50°S) is the strongest sink providing about 60% of the total global ocean uptake, whereas the Pacific (north of 50°S) is nearly neutral. The Indian and Southern Oceans contribute about 20% each to the global uptake flux. The uptake flux by the North Pacific is similar to that by the South Pacific, whereas the North Atlantic takes up 0.45 to 0.6 Gt-C·yr−1 more CO2 than the South Atlantic and Southern Ocean combined. This is comparable to the preindustrial oceanic transport of CO2 from the North to South Atlantic estimated by Broecker and Peng (11). Because the results of this study differ significantly from the northern ocean uptake of CO2 used in the analysis of the atmospheric and oceanic data by Tans et al. (1), a global analysis of the CO2 and carbon isotope data in the ocean, atmosphere, and biosphere must be made to evaluate their mutual coherence.
This study has benefited from observations made by many international investigators. We thank the following scientists for their contributions: W. S. Broecker, A. W. Dickson, R. H. Gammon, S. S. Jacobs, Walker Smith, H. Ducklow, W. M. Smethie, D. Martinson, P. Schlosser, J. Sarmiento, D. Wallace, E. Garvey, N. R. Bates, A. H. Knap, T. D. Foster, C. S. Wong, C. D. Keeling, L. Merlivat, C. LeQuere, V. Garcon, C. Provost, W. Roether, H. Y. Inoue, K. Fushimi, A. Watson and J. E. Robertson. We also thank members of our technical staff who ran instruments at sea and on land and processed data in our respective laboratories: J. Goddard, S. Rubin, R. Esmay, F. A. van Woy, P. K. Salameh, P. P. Murphy, K. C. Kelly, L. S. Waterman, Matt Steckley and David Ho. This work was supported by a number of grants from the National Science Foundation, the U.S. Department of Energy, and the National Oceanic and Atmospheric Administration to T.T. at the Lamont–Doherty Earth Observatory and to R.F.W. at the Scripps Institution of Oceanography. R.A.F. and R.H.W. have been supported by the Climate and Global Change Program of the National Oceanic and Atmospheric Administration. We gratefully acknowledge the support and encouragement received from these agencies. This is contribution 5573 of the Lamont–Doherty Earth Observatory.
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