Chapter 7: APPLICATIONS
Figure 7-20 shows a summary of gas transfer results over the ocean. The
theoretical dependencies correspond to the sum of an interfacial component
(renewal model described in Section 2.3.2), gas transfer by wave-breaking
turbulent patches described by (7.25), and a bubble-mediated component
described by equation (7.28).
Both theoretical relationships and field data are color-coded. Blue color
indicates low solubility gases (SF 6 and
3 He); red color indicates high
solubility gases (Rn and CO 2 ). Under high wind speed conditions, the
theoretical dependencies are quite different depending on gas solubility but
are consistent with the available data.
The model shown in Figure 7-20 is able to explain the difference
between the Wanninkhof (1992) and Liss and Merlivat (1986) empirical
parameterizations for gas transfer velocity. The former was derived from
dual tracer data (using low solubility gases SF 6 and
3 He), while the latter was
derived from a laboratory experiment with a highly soluble gas (CO 2 ).
7.5.5 Fine thermohaline structure and gas exchange
The near-surface layer is not always well mixed. Appreciable
temperature, and salinity differences can develop in the upper few meters of
the ocean caused by the diurnal cycle and precipitation effects under low
wind speed conditions. The temperature difference across the diurnal
thermocline can reach as much as several
o C and the salinity difference
across the rain-formed halocline as much as several psu (see Chapter 4 for
details). Dependence of the CO 2 solubility on temperature suggests that
diurnal warming may shift the concentration difference between atmosphere
and uppermost layer of the sea towards lower or negative CO 2 uptake by the
ocean.
In addition, the presence of the diurnal thermocline (or a rain-formed
halocline) may cause depletion or accumulation of gas in the near-surface
layer. This effect, however, is not expected to change appreciably the CO 2
concentration in uppermost layer of the ocean due to the large buffer
capacity of the carbonate system in seawater (but may be significant for
other gases like O 2 ).
In the commonly used parameterization formulation for the net air-sea
gas flux (1.45), the gas partial pressure w
p in the bulk of water is usually
taken from the ship’s intake or from CTD measurements at roughly 3-5 m
depth. Thermosalinograph or CTD data from the same depth are also usually
used to calculate the gas solubility, S P , which is a function of temperature
and salinity. As a result, the large diurnal warming events, which are
localized in the upper few meters of the ocean, are often undetected during
shipboard surveys and unaccounted for when quantifying air-sea gas fluxes.
489
Figure 7-20 shows a summary of gas transfer results over the ocean. The
theoretical dependencies correspond to the sum of an interfacial component
(renewal model described in Section 2.3.2), gas transfer by wave-breaking
turbulent patches described by (7.25), and a bubble-mediated component
described by equation (7.28).
Both theoretical relationships and field data are color-coded. Blue color
indicates low solubility gases (SF 6 and
3 He); red color indicates high
solubility gases (Rn and CO 2 ). Under high wind speed conditions, the
theoretical dependencies are quite different depending on gas solubility but
are consistent with the available data.
The model shown in Figure 7-20 is able to explain the difference
between the Wanninkhof (1992) and Liss and Merlivat (1986) empirical
parameterizations for gas transfer velocity. The former was derived from
dual tracer data (using low solubility gases SF 6 and
3 He), while the latter was
derived from a laboratory experiment with a highly soluble gas (CO 2 ).
7.5.5 Fine thermohaline structure and gas exchange
The near-surface layer is not always well mixed. Appreciable
temperature, and salinity differences can develop in the upper few meters of
the ocean caused by the diurnal cycle and precipitation effects under low
wind speed conditions. The temperature difference across the diurnal
thermocline can reach as much as several
o C and the salinity difference
across the rain-formed halocline as much as several psu (see Chapter 4 for
details). Dependence of the CO 2 solubility on temperature suggests that
diurnal warming may shift the concentration difference between atmosphere
and uppermost layer of the sea towards lower or negative CO 2 uptake by the
ocean.
In addition, the presence of the diurnal thermocline (or a rain-formed
halocline) may cause depletion or accumulation of gas in the near-surface
layer. This effect, however, is not expected to change appreciably the CO 2
concentration in uppermost layer of the ocean due to the large buffer
capacity of the carbonate system in seawater (but may be significant for
other gases like O 2 ).
In the commonly used parameterization formulation for the net air-sea
gas flux (1.45), the gas partial pressure w
p in the bulk of water is usually
taken from the ship’s intake or from CTD measurements at roughly 3-5 m
depth. Thermosalinograph or CTD data from the same depth are also usually
used to calculate the gas solubility, S P , which is a function of temperature
and salinity. As a result, the large diurnal warming events, which are
localized in the upper few meters of the ocean, are often undetected during
shipboard surveys and unaccounted for when quantifying air-sea gas fluxes.
489