Exchange of CO2 under low turbulence conditions and high pH can be enhanced by hydration reactions of CO, with hydroxide ions and water molecules in the boundary layer. A series of field experiments was performed on several lakes, including alkaline closed-basin lakes, using enclosures (helmets) to study the enhancement process in nature. In addition, the enhancement of CO, exchange was studied in laboratory experiments with freshwater and seawater. The results of the experiments are compared with published theoretical calculations. Within the experimental uncertainties and shortcomings of the chemical enhancement models, reasonable agreement was observed between experimental and theoretical results for seawater. The experiments indicate, in accordance with theory, that chemical enhancement has a minor effect on air-sea gas exchange of CO, under average oceanic turbulence conditions. However, for the equatorial CO, source regions, with high temperatures and low winds, the calculated CO, enhancement amounts to 4-8% of the total exchange. The observations on lakes show poorer agreement with models, which is attributed to experimental uncertainty and poor characterization of the chemistry of the lake waters. The experiments show that chemical enhancement of CO, is ubiquitous for the alkaline closed-basin lakes with enhancements of up to a factor of three.Air-water gas transfer for slightly soluble gases is retarded in the frst several hundred microns of the liquid (Liss 1983). Slightly soluble gases arc defined as those that have Ostwald solubility coefficients (0) of < 10, where p is the volum,e of gas at temperature T and partial pressure p dissolved per unit volume of water (thus ,G' is dimensionless). In the simplest gas transfer model-the stagnant boundary layer model-transfer through the boundary layer occurs by molecular diffusion while above and below the liquid boundary layer transport occurs by much faster turbulent processes (Liss and Slater 1974). The gas transfer velocity (k)(cm h-l) through the stagnant boundary layer can thus be defined as k = 3,600 D z-l.(1) D is the molecular diffusion coefficient of the gas (cm* s-l), 3,600 is the conversion from hours to seconds, and z is the stagnant boundary layer thickness (cm) (Fig. 1). Chemically reactive gases can bypass the rate-limiting molecular diffusion step by reacting in the boundary layer. Although theory and several experimental studies indicate that the stagnant film model is conceptually incorrect, the principle of chemical enhancement remains unchanged, irrespective of gas transfer model. CO2 gas can undergo two hydration reactions in water:CO2 + OH-= HC03-.i AcknowledgmentsWe thank Rachel Oxburgh and Rob No11 for assisting in the experiments on the closed-basin lakes. Kitack Lee performed the alkalinity analyses of the seawater samples. Cathy Costa and Ed Harrison supplied wind-speed data for the equatorial Pacific enhancement exercise. Suggestions by two anonymous reviewers and Bernard Boudreau were most helpful.This research was suppo...
We present experimental results that show that the kinetic isotopic fractionation during gas exchange is 0.9972 ± 0.0002 for oxygen, 0.9992 ± 0.0002 for methane, 0.9987 ± 0.0001 for nitrogen and 0.982 ± 0.002 for hydrogen, and that the equilibrium fractionation between water and gas phases is 1.037 for hydrogen. We show that the isotopic fractionation during gas transfer for these gases is not equal to the square root of their reduced mass in water, as would be predicted by an extension of the kinetic theory of ideal gases to dissolved gases. The use of isotopes as tracers of biogeochemical gases requires knowledge of the fractionation factor for air‐water gas transfer; there have been few direct measurements of these factors.
Distributions of oxygen, argon, nitrogen, and radon in the upper ocean of the subarctic Pacific distinguish the fluxes controlling the oxygen mass balance during the summers of 1987 and 1988. The difference between the net O2 flux (in mmol m−2 d−1) to the atmosphere via gas exchange (32) and the integrated decrease with time (−14) is balanced by biological production (13‐17), air injection by bubble entrainment (5), and O2 flux to the thermocline −(0‐4). Nitrogen/argon and oxygen/argon ratios reveal that ˜15% of the oxygen supersaturation in summer is produced by air injection and ˜40% by biological production, with the rest induced by surface water warming. Our estimate of biologically induced oxygen production when translated stoichiometrically to nitrogen uptake agrees to within error estimates with both the particulate and dissolved nitrogen mass balances for the upper ocean determined in the SUPER program during the same time period. The oxygen mass balance requires a net carbon production in the euphotic zone of ˜140 mg C m−2 d−1 (PQ=1.5), which is 20–30% of the level of 14C primary production determined by SUPER investigators.
The 18O/16O of dissolved oxygen was measured in the upper ocean of the subarctic Pacific in 1988. In May and August, at stations Papa (50°N, 145°W) and R (53°N, 145°W), the mean δ18O in the mixed layer was 23.84±0.20 and 24.00±0.24 ‰ (versus standard mean ocean water) consistently more depleted than atmospheric saturation levels by about 0.4 ‰. This relative depletion is caused by input of photosynthetically produced O2. A value for the isotopic fractionation effect during respiration (αT) of 0.978±0.006 was determined from the time rate of change of the concentration and δ18O of O2 in the mixed layer measured during August 1988. Below the mixed layer (100–280 m) the O2 concentration decreased with a corresponding increase in δ18O. Model derived values for αT over this depth region ranged from 0.980 to 0.988 and depended on the mixing model. The difference between αT determinations for the surface layer versus upper thermocline likely results from mixing model inaccuracies or different isotope fractionation effects during plankton and bacterial respiration. If the calculated mixed layer αT values apply oceanwide, then photosynthesis and respiration by the marine biota have a similar effect to land plants in maintaining the δ18O of atmospheric O2 at 23.5 ‰.
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