Fluxes of Sensible Heat, Latent Heat, Impulse, and Atmospheric Water Vapor over the North Atlantic from the EOS Aqua AMSR-E Radiometer

Abstract

Possibilities of using the EOS Aqua AMSR-E MCW radiometer for analysis of the long-term dynamics of the Gulf Stream heat regime are analyzed. The SOA brightness temperature, total water vapor content of the atmosphere, and surface heat, water, and impulse fluxes tracks along the Gulf Stream and North Atlantic current are under consideration. Their temporal and spatial variability is analyzed in the time periods preceding intensive oil spills in the Gulf of Mexico in April 2010 year and following periods. Possibilities of using the SOA brightness temperature measured by various channels of the AMSR-E radiometer as the immediate characteristic of the ocean-atmosphere heat interaction and their seasonal variability are shown.

Method for Computing the Monthly Mean Fluxes of Heat, Moisture, and Impulse from the Radiometer AMSR-E

Initial relationships and approximations. At present, a more realistic scenario for estimating the surface fluxes of sensible heat, latent heat, and impulse at the ocean–atmosphere boundary are the bulk formulas cited in Chap. 1. The important points in this analysis of ocean surface temperature (OST), near-surface air temperature, humidity, and wind speed can be determined from the EOS Aqua AMSR-E radiometer.

To estimate the relative near-surface air humidity $q_{\text{a}}$, we used the following relationships, where the data of measurements of the SOA brightness temperature in all twelve AMSR-E radiometric channels were used:

$$ \begin{aligned} q_{\text{a}} & = a_{0} + a_{1} T_{{6{\text{V}}}} + a_{2} T_{{6{\text{H}}}} + a_{3} T_{{10{\text{V}}}} + a_{4} T_{{10{\text{H}}}} + a_{5} T_{{18{\text{V}}}} + a_{6} T_{{18{\text{H}}}} \\ & \quad + a_{7} T_{{22{\text{V}}}} + a_{8} T_{{22{\text{H}}}} + a_{9} T_{{36{\text{V}}}} + a_{10} T_{{36{\text{H}}}} + a_{11} T_{{89{\text{V}}}} + a_{12}^{{}} T_{{89{\text{H}}}} \\ \end{aligned} $$

(8.1)

Here, the digital indexes of T denote the frequency of the radiometric channel (in GHz); the symbols V and H indicate vertical and horizontal polarization, respectively.

To estimate the near-surface air temperature, the parameterization is used (Algorithm, HOAPS 2011; Andersson et al. 2010):

$$ T_{\text{a}} = 1.03T_{\text{S}} - 1.32 $$

(8.2)

Computation of the saturated relative humidity e ₀ was performed using Magnus’ formula:

$$ e_{0} = 6.1078 \cdot \exp \left[ {\frac{{ 1 7. 2 6 9 3 8 8 2\cdot \left( {{\text{T}}_{\text{S}} - 273.16} \right)}}{{T_{\text{S}} - 3 5. 8 6}}} \right] $$

(8.3)

The following relationships were used to determine the saturated near-surface air humidity over a salty water surface proper for the ocean (q _s), proper for the ocean surface:

$$ e_{{0{\text{S}}}} = 0.98e_{0} $$

(8.4)

$q_{\text{S}} = 0.622\frac{{e_{{{\text{S}}0}} }}{{p - 0.378e_{{{\text{S}}0}} }}$, at the standard near-surface atmospheric pressure.

The main problem of the Global Aerodynamic Method is the alternatives for determining the coefficients in these formulas. As a rule, the values of the heat and moisture exchange in the SOA interface vary essentially from $C_{\text{H}}$ = (1−2) × 10⁻³, $C_{\text{E}}$ = (1.0–1.7) × 10⁻³ (Lappo et al. 1990).

Figure 8.12 illustrates the dependence of the number of Schmidt (C _Т) from the near-surface wind speed on the coefficient $C_{\text{H}}$ for the two parameterizations, which denote go-no-go scattering of this parameter in many other known variants (cited in Lappo et al. 1990, resting upon the works of Garrat 1977 and Kondo 1975).

Fig. 8.12

Dependence of the Schmidt index on the near-surface wind speed due to Garrat and 2 (Condo)

Figure 8.12 shows that maximal distinctions between various parameterizations of heat and moisture exchanges in the SOA interface are apparent in a weak weather force under wind speed to 3 m s⁻¹. Panin (1987), Panin and Krivitskii (1992) provided useful relationships for sensible and latent fluxes in these conditions:

$$ H = A\uprho_{\text{a}} c_{\text{p}} \left( {T_{\text{S}} - T_{\text{a}} } \right)^{4/3} \left[ {\alpha gk_{\text{T}}^{2} \nu^{ - 1} (1 + b/{\text{Bo}})} \right]^{1/3} ; $$

(8.5)

$$ L_{\text{E}} = A\uprho_{\text{a}} L_{\text{S}} \left( {q_{\text{S}} - q_{\text{a}} } \right)^{4/3} \left[ {\beta gk_{\text{q}}^{2} \nu^{ - 1} (1 + {\text{Bo}}/b)} \right]^{1/3} . $$

(8.6)

Here, $A$ = 0.15, $\beta$ = 0.61, L _S = 25.04 × 10⁵ J kg⁻¹ is the specific evaporation heat, g = 9.81 m s^−2,, α is the coefficient of the air heat expanding (b ≈ 0.073), $\nu$ is the air kinematic viscosity (thickness); ${\text{Bo}} = \frac{H}{{L_{\text{E}} }}$ is the Bowen number; and $k_{\text{T}} ,k_{\text{q}}$ are the kinematical coefficients of the molecular diffusion of heat and water vapor exchanges.

For wind speed exceeding 3 m s⁻¹, the values of the heat exchange coefficients were taken as $C_{\text{H}}$ = 0.0012 and $C_{\text{E}}$ = 0.0011. For evaluation of the resistance coefficient, we summarized the data of Panin (1987), Panin and Krivitskii (1992) and Repina (2007) in the form of piecewise linear dependence $C_{\text{D}}$ = a + b (V − с) in the following conditions:

$$ \begin{aligned} & {\text{For}}\,V < 3\,{\text{m}}^{ - 1} ,C_{\text{D}} = 0.001\,{\text{does}}\,{\text{not}}\,{\text{depend}}\,{\text{on}}\,V \\ & {\text{For}}\,V = 3 - 12.5\,{\text{m}}\,{\text{s}}^{ - 1} ,a = 1,\,b = 0.0706,\,{\text{and}}\,c = 3 \\ & {\text{For}}\,V > 12.5\,{\text{m}}\,{\text{s}}^{ - 1} ,a = 1.6,\,b = 0.02286,\,{\text{and}}\,c = 12.5 \\ \end{aligned} $$