AMSR Ocean Algorithm
Frank J. Wentz
Remote Sensing Systems

1. Overview and Background Information
1.1. Introduction
With the advent of well-calibrated satellite microwave radiometers, it is now possible to obtain long time series of geophysical parameters that are important for studying the global hydrologic cycle and the Earth's radiationbudget. Over the world's oceans, these radiometers simultaneously measure profiles of air temperature and thethree phases of atmospheric water (vapor, liquid, and ice). In addition, surface parameters such as the nearsurface wind speed, the sea-surface temperature, and the sea ice type and concentration can be retrieved. A wide
variety of hydrological and radiative processes can be studied with these measurements, including air-sea and airice interactions (i.e., the latent and sens ible heat fluxes, fresh water flux, and surface stress) and the effect ofclouds on radiative fluxes. The microwave radiometer is truly a unique and valuable tool for studying our planet.
This Algorithm Theoretical Basis Document (ATBD) focuses on the Advanced Microwave Scanning Radiometer (AMSR) that is scheduled to fly in December 2000 on the NASA EOS-PM1 platform. AMSR will measure the Earth’s radiation over the spectral range from 7 to 90 GHz. Over the world’s oceans, it will be possible toretrieve the four important geophysical parameters listed in Table 1. The rms accuracies given in Table 1 comefrom past investigations and on-going simulations that will be discussed. Rainfall can also be retrieved, which is
discussed in a separate AMSR ATBD.
We are confident that the expected retrieval accuracies for wind, vapor, and cloud will be achieved. The Special Sensor Microwave Image (SSM/I) and the TRMM microwave imager (TMI) have already demonstrated thatthese accuracies can be obtained. The AMSR wind retrievals will probably be more accurate than that of SSM/I
and less affected by atmospheric moisture. A comparison between sea surface temperatures (SST) from TMI withbuoy measurements indicate an rms accuracy between 0.5 and 0.7 K. One should keep in mind that part of theerror arises from the temporal and spatial mismatch between the buoy measurement and the 50 km satellite footprint. Furthermore, the satellite is measuring the temperature at the surface the ocean (about 1 mm deep) whereas
the buoy is measuring the bulk temperature near 1 m below the surface. There are still some concerns with regards to the sea-surface temperature retrieval, which are discussed in Section 1.5.
This document is version 2 of the AMSR Ocean Algorithm ATBD. The primary difference between this version
and the earlier version is that the emissivity model for the 10.7 GHz has been updated using data from TMI. In
addition, there are several small updates to the radiative transfer model (RTM).
Table 1. Expected Retrieval Accuracy for the Ocean Products
Geophysical Parameter
Rms Accuracy
Sea-surface temperature TS
0.5 K
Near-surface wind speed W
1.0 m/s
Vertically integrated (i.e., columnar) water vapor V
1.0 mm
Vertically integrated cloud liquid water L
0.02 mm1.2. Objectives of Investigation
There are two major objectives of this investigation. The first is to develop an ocean retrieval algorithm thatwill retrieve TS, W, V, and L to the accuracies specified in Table 1. These products will be of great value to the
Earth science community. The second objective is to improve the radiative transfer model (RTM) for the ocean
surface and non-raining atmosphere. The 6.9 and 10.7 GHz channels on AMSR will provide new information on
the RTM at low frequencies. Experience has shown that these two objectives are closely linked. A better understanding of the RTM leads to more accurate retrievals. A better understanding of the RTM also leads to new remote sensing techniques such as using radiometers to measure the ocean wind vector.
1.3. Approach to Algorithm Development
Radiative transfer theory provides the relationship between the Earth’s brightness temperature TB (K) as
measured by AMSR and the geophysical parameters TS, W, V, and L. This ATBD addresses the inversion problem of finding TS, W, V, and L given TB. We place a great deal of emphasis on developing a highly accurate
RTM. Most of our AMSR work thus far has been the development and refinement of the RTM. This work is now
completed, and Section 2 describes the RTM in considerable detail.

The importance of the RTM is underscored by the fact that AMSR frequency, polarization, and incidence angle selection is not the same as previous satellite radiometers. Table 2 compares AMSR with other radiometersystems. Albeit some of the differences are small, they are still significant enough to preclude developing AMSR
algorithms by simply using existing radiometer measurements. The differences in frequencies and incidence angle must be taken into account when developing AMSR algorithms.
2. Geophysical Model for the Ocean and Atmosphere2.1. Introduction
The key component of the ocean parameter retrieval algorithm is the geophysical model for the ocean andatmosphere. It is this model that relates the observed brightness temperature (TB) to the relevant geophysical parameters. In remote sensing, the specification of the geophysical model is sometimes referred to as the forward
problem in contrast to the inverse problem of inverting the model to retrieve parameters. An accurate specific ation of the geophysical model is the crucial first step in developing the retrieval algorithm.
2.2. Radiative Transfer Equation
We begin by deriving the radiative transfer model for the atmosphere bounded on the bottom by the Earth’s
surface and on the top by cold space. The derivation is greatly simplified by using the absorption-emission approximation in which radiative scattering from large rain drops and ice particles is not included. Over the spectral
range from 6 to 37 GHz, the absorption-emission approximation is valid for clear and cloudy skies and for lightrain up to about 2 mm/h. The results of Wentz and Spencer [1997] indicate that only 3% of the SSM/I observations over the oceans viewed rain rates exceeding 2 mm/h. Thus, the absorption-emission model to be presented
will be applicable to about 97% of the AMSR ocean observations.
In the microwave region, the radiative transfer equation is generally given in terms of the radiation brightness
temperature (TB), rather than radiation intensity. So we first give a brief discussion on the relationship between
radiation intensity and TB. Let Pλ denote the power within the wavelength range dλ, coming from a surface area
dA, and propagating into the solid angle dΩ. The specific intensity of radiation Iλ is then defined by
Pλ = I λ cos θi d λ dA d Ω(1)
The specific intensity is in units of erg/s-cm3 -steradian. The angle θi is the incidence angle defined as the angle
between the surface normal and the propagation direction. For a black body, Iλ is given by Planck’s law to be[Reif, 1965]

Iλ =

2hc 2
λ5 [exp (hc λkT ) − 1](2)

where c is the speed of light (2.998×1010 cm/s), h is Planck’s constant (6.626×10−27 erg-s), k is Boltzmann’s constant (1.381×10−16 erg/K), λ (cm) is the radiation wavelength, and T (K) is the temperature of the black body.
Consider a surface that is emitting radiation with a specific intensity Iλ. Then the brightness temperature TB for
this surface is defined as the temperature at which a black body would emit the radiation having specific intensity
Iλ. In the microwave region, the exponent in (2) is small compared to unity, and (2) can be easily inverted to giveT B in terms of Iν.
λ4 I λ
TB =
(3)
2 kc
This approximation is the well known Rayleigh Jeans approximation for λ >> hc/kT.
In terms of TB, the differential equation governing the radiative transfer through the atmosphere is∂T B= −α ( s ) T B ( s ) − T ( s )(4)∂swhere the variable s is the distance along some specified path through the atmosphere. The terms α(s) and T(s)are the absorption coefficient and the atmospheric temperature at position s. Equation (4) is simply stating thatthe change in TB is due to (1) the absorption of radiation arriving at s and (2) to emission of radiation emanatingfrom s. We let s = 0 denote the Earth’s surface, and let s = S denote the top of the atmosphere (i.e., the elevation
above which α(s) is essentially zero).
Two boundary conditions that correspond to the Earth’s surface at the bottom and cold space at the top are
applied to equation (4). The surface boundary conditions states that the upwelling brightness temperature at thesurface TB↑ is the sum of the direct surface emission and the downwelling radiation that is scattered upward by therough surface [Peake, 1959].

π 2

2π
sec θiT B ↑ (k i ,0) = E( k i )TS +sin θs dθs ∫ dϕs TB ↓ (k s ,0)[σo , c (k s , k i ) + σo ,× (k s , k i )]
4π ∫0
0

(5)

where the first TB argument denotes the propagation direction of the radiation and the second argument denotes
the path length s. The unit propagation vectors k i and k s denote the direction of the upwelling and downwellingradiation, respectively. In terms of polar angles in a coordinate system having the z-axis normal to the Earth’s
surface, these propagation vectors are given by
k i = cos ϕ i sin θ i , sin ϕi sin θ i , cos θ i(6a)k s = − cos ϕs sin θ s , sin ϕs sin θ s , cos θs(6b)
The first term in (5) is the emission from the surface, which is the product of the surface temperature TS and thesurface emissivity E(k i ). The second term is the integral of downwelling radiation TB↓(k s) that is scattered in direction ki . The integral is over the 2π steradian of the upper hemisphere. The rough surface scattering is characterized by the bistatic normalized radar cross sections (NRCS) σo,c(θs ,θi) and σ o,×(θs,θi ). These cross sections
specify what fraction of power coming from ks is scattered into k i. The subscripts c and × denote co-polarization(i.e., incoming and outgoing polarization are the same) and cross-polarization (i.e., incoming and outgoing polarizations are orthogonal), respectively. The cross sections also determine the surface reflectivity R(k i) via the following integral.
π 2

2π
sec θiR(k i ) =sin θs dθ s ∫ dϕs [σo ,c (k s , k i ) + σ o ,× (k s , k i )]
4π ∫0
0

(7)

The surface emissivity E(k i ) is given by Kirchhoff’s law to beE(k i ) = 1 − R(k i )(8)
It is important to note that equations (5) an (7) provide the link between passive microwave radiometry and active
microwave scatterometry. The scatterometer measures the radar backscatter coefficient, which is simply σ o,c(k i ,k i ).
The upper boundary condition for cold space isT B ↓ (k s , S ) = TC
(9)This simply states that the radiation coming from cold space is isotropic and has a magnitude of TC = 2.7 K.
The differential equation (4) is readily solved by integrating and applying the two boundary conditions (5)and (9). The result for the upwelling brightness temperature at the top of the atmosphere (i.e., the value observedby Earth-orbiting satellites) is
TB ↑ (k i , S ) = TBU + τ[ETS + TBΩ ](10)
where TBU is the contribution of the upwelling atmospheric emission, τ is the total transmittance from the surface
to the top of the atmosphere, E is the Earth surface emissivity given by (8), and TBΩ is the surface scattering integral in (5). The atmospheric terms can be expressed in terms of the transmittance function τ(s 1 ,s 2 ) s2τ(s1 , s 2 ) = exp  − ∫ ds α( s) 
 s

 1


(11)

which is the transmittance between points s1 and s2 along the propagation path k s or ki. The total transmittance τin (10) is given byτ = τ(0, S )(12)and the upwelling and downwelling atmosphere emissions are given byS

T BU = ∫ ds α( s ) T ( s ) τ( s , S )

(13a)

0
S

T BD = ∫ ds α( s ) T ( s ) τ (0, s )

(13b)

0

The sky radiation scattering integral is
π 22π
sec θi
TB Ω =sin θs dθs ∫ dϕs (TBD + τTC )[σo , c (k s , k i ) + σo ,× (k s , k i )]
4π ∫0
0

(14)

Thus, given the temperature TS and absorption coefficient α at all points in the atmosphere and given the surface bistatic cross sections, T B can be rigorously calculated. However, in practice, the 3-dimensional specificationof TS and α over the entire volume of the atmosphere is not feasible, and to simplify the problem, the assumptionof horizontal uniformity is made. That is to say, the absorption is assumed to only be a function of the altitude habove the Earth’s surface, i.e., α(s) = α(h). To change the integration variable from ds to dh, we note that for the
spherical Earth
∂s
1+δ
=
(15)
∂h
cos 2 θ + δ(2 + δ)
where θ is either θi or θs , δ = h/RE , and RE is the radius of the Earth. In the troposphere δ << 1, and an excellent
approximation for θ < 60° is,∂s= sec θ(16)∂hWith this approximation and the assumption of horizontal uniformity, the above equations reduce to the following
expressions.
h2τ(h1 , h2 ,θ ) = exp − sec θ ∫ dh α( h) 


h1


τ = τ(0, H ,θi )

(17)
(18)

H

T BU = secθi ∫ dh α( h) T ( h) τ (h, H ,θi )

(19a)

0

H

TBD = sec θ s ∫ dh α( h) T ( h) τ( 0, h,θs )

(19b)

0

Thus, the brightness temperature computation now only requires the vertical profiles of T(h) and α(h) along withthe surface cross sections. The following two sections discuss the atmospheric model for α(h) and the sea-surface
model for the cross sections, respectively. In closing, we note that the AMSR incidence angle is 55° and hence
approximation (16) is quite valid, with one exception. In the scattering integral, θs goes out to 90°, and in thiscase we use (15) to evaluate the integral.
2.3. Model for the Atmosphere
In the microwave spectrum below 100 GHz, atmospheric absorption is due to three components: oxygen, water vapor, and liquid water in the form of clouds and rain [Waters, 1976]. The sum of these three components
gives the total absorption coefficient (napers/cm).
α( h ) = α O ( h ) + α V ( h ) + α L ( h )(20)
Numerous investigators have studied the dependence of the oxygen and water vapor coefficients on frequency ν(GHz), temperature T (K), pressure P (mb), and water vapor density ρV (g/cm3 ) [Becker and Autler, 1946; Rozenkranz, 1975; Waters, 1976; Liebe, 1985]. To specify α O and α V as a function of (ν,T,P,ρV) we use the Liebe[1985] expressions with one modification. The self-broadening component of the water vapor continuum is reduced by a factor of 0.52 (see below). The liquid water coefficient αL comes directly from the Rayleigh approximation to Mie scattering and is a function of T and the liquid water density ρL (g/cm2 ) (see below). Figure 3
shows the total atmospheric absorption for each component. Results for three water vapor cases (10, 30, and 60mm) are shown. The cloud water content is 0.2 mm. This corresponds to a moderately heavy non-raining cloudlayer.
Let AI denote the vertically integrated absorption coefficient.
H

AI = ∫ dh αI ( h)

(21)

0

where h is the height (cm) above the Earth’s surface and subscript I equals O, V, or L. Equations (17) and (18)then give the total transmittance to beτ = exp [− secθi ( AO + AV + AL )](22)
Assuming for the moment that the atmospheric temperature is constant, i.e., T(h) = T, then the integrals in equations (19) can be exactly evaluated in closed form to yield

TBU = TBD = (1 − τ )T(23)
In reality, the atmospheric temperature does vary with h, typically decreasing at a lapse rate of about -5.5 C/km in
the lower to mid troposphere. In view of (23), we find it convenient to parameterize the atmospheric model interms of the following upwelling and downwelling effective air temperatures:
T U = TB U / (1 − τ )(24a)
T D = TBD / (1 − τ )(24b)These effective temperatures are indicative of the air temperature averaged over the lower to mid troposphere.
Note that in the absence of significant rain, TU and TD are very similar in value, with TU being 1 to 2 K colder.
100.

Total Atmosphere Attenuation (naper)

10.

1.0

0.1

0.01

LegendOxygenWater Vapor, 10 mm
Water Vapor, 30 mm
Water Vapor, 60 mm
Cloud, 0.20 mm

0.001

0.0001

0

10

20

30

40

50

60

70

80

90

100

Frequency (GHz)
Fig. 3. The atmospheric absorption spectrum for oxygen, water vapor, and cloud water. Results for three water vapor cases (10,30, and 60 mm) are shown. The cloud water content is 0.2 mm which corresponds to a moderately heavy non-raining cloud layer.

In view of the above equations, one sees that the atmospheric model can be parameterized in terms of the following 5 parameters:
1. Upwelling effective temperature TU
2. Downwelling effective temperature TD
3. Vertically integrated oxygen absorption AO
4. Vertically integrated water vapor absorption AV
5. Vertically integrated liquid water absorption AL
To study the properties of the first four parameters, we use a large set of 42,195 radiosonde flights launched fromsmall islands [Wentz, 1997]. These radiosonde reports provide air temperature T(h), air pressure p(h), and watervapor density ρV(h) at a number of levels in the troposphere. From these data, the coefficients α O and α V are
computed from the Liebe [1985] expressions, except that the water vapor continuum term is modified as discussed
in the next paragraph. Performing the numerical integrations as indicated above, TU , TD, AO , and AV are found foreach radiosonde flight. In addition, the vertically integrated water vapor V is also computed.
H

V = 10 ∫ dh ρV ( h)

(25)

0

where ρV(h) is in units of g/cm3 , and the leading factor of 10 converts from g/cm2 to mm.
Wentz [1997] computed AV directly from collocated SSM/I and radiosonde observations. At 19, 22, and 37
GHz, the Liebe AV was found to be 4%, 3%, and 20% higher than the SSM/I-derived value, respectively. To
quote Liebe [1985]: ‘Water vapor continuum absorption has been a major source of uncertainty in predicting millimeter wave attenuation rates, especially in the window ranges.’ The frequency of 37 GHz is in a water vapor
window and is most affected by the continuum. It should be noted that Liebe also needed to rely on combined
radiometer-radiosonde measurements to infer the continuum in the 6 to 37 GHz region. Liebe’s data set in thisspectral region is rather limited and does not contain any 37 GHz observations. We believe the SSM/I method ofderiving AV is more accurate than Liebe’s method, and hence adjust the Liebe [1985] water vapor spectrum so that
it will agree with the SSM/I results. We find that very good agreement is obtained by reducing the selfbroadening component of the water vapor continuum by a factor of 0.52. After this adjustment, the agreement at
all three frequencies is within ± 1%.
Figure 4 shows the T D values computed from the 42,195 radiosondes plotted versus V. Three frequencies areshown (19, 22, and 37 GHz), and the curves are quite similar. The solid lines in the figure show equation (26),and vertical bars show the ± one standard deviation of TD derived from the radiosondes. For low to moderate values of V (0 to 40 mm), TD increases with V, and above 40 mm, TD reaches a relatively constant value of 287 K.
The TU versus V curves (not shown) are very similar except that TU is 1 to 2 K colder. The following least-square
regressions are found to be a good approximation of the TD , T U versus V relationship:(26a)
T U = TD + b 6 + b 7 V
(26b)
where
T V = 273 .16 + 0. 8337 V − 3. 029 × 10 −5 V 3 . 3 3
V ≤ 48(27a)
T V = 301.16
V > 48(27b)andς(TS − TV ) = 1.05 ⋅ TS − TV ⋅ 1 −(TS − TV ) 2

1200
ς (TS − TV ) = sign(TS − TV ) ⋅ 14K| TS − TV | ≤ 20K(27c)| TS − TV | > 20K(27d)

V is in units of millimeters and all temperatures are in units of Kelvin. When evaluating (26a), the expression islinearly extrapolated when V is greater than 58 mm. We have included a small additional term that is a functionof the difference between the sea-surface temperature TS and TV, which represents the sea-surface temperaturethat is typical for water vapor V. The term ς ( TS − TV ) accounts for the fact that the effective air temperature istypically higher (lower) for the case of unusually warm (cold) water. The TV versus V relationship was obtainedby regressing the climatology sea-surface temperature at the radiosonde site to V derived from the radiosondes.
Over the full range of V, the rms error in approximation (26) is typically about 3 K. Table 4 gives the b0 throughb7 coefficients for all 8 AMSR frequencies.
The vertically integrated oxygen absorption AO is nearly constant over the globe, with a small dependence on the
air temperature. We find the following expression to be a very good approximation for AO :AO = aO 1 + aO 2 (TD − 270 )(28)
Table 4 gives the aO coefficients for the 8 AMSR frequencies, and Table 5 gives the rms error in this approximation for the 8 frequencies. At 23.8 GHz and below, the error is negligible, being 0.0003 napers or less. At 36.5
GHz, the error is still quite small, being 0.0008 napers. Note that 0.001 napers roughly corresponds to a T B errorof 0.5 K. For the 50.3 and 52.8 GHz oxygen band channels, the error is considerably larger, but (28) is not used
for the oxygen band channels. Rather the oxygen band channels can be used to retrieve TD .

Table 4. Model Coefficients for the Atmosphere
Freq. (GHz)

6.93E+010.65E+0

18.70E+023.80E+036.50E+050.30E+052.80E+0

89.00E+0b0 (K)

239.50E+0239.51E+0240.24E+0241.69E+0239.45E+0242.10E+0245.87E+0

242.58E+0

b1 (K mm −1 )b2 (K mm−2 )213.92E−2−460.60E−4

225.19E−2−446.86E−4

298.88E−2−725.93E−4

310.32E−2−814.29E−4

254.41E−2−512.84E−4

229.17E−2−508.05E−4

250.61E−2−627.89E−4

302.33E−2−749.76E−4

b3 (K mm−3 )457.11E−6

391.82E−6

814.50E−6

998.93E−6

452.02E−6

536.90E−6

759.62E−6

880.66E−6

b4 (K mm−4 )−16.84E−7−12.20E−7−36.07E−7−48.37E−7−14.36E−7−22.07E−7−36.06E−7−40.88E−7

b5

0.50E+0

0.54E+0

0.61E+0

0.20E+0

0.58E+0

0.52E+00.53E+0

0.62E+0

b6 (K)−0.11E+0−0.12E+0−0.16E+0−0.20E+0−0.57E+0−4.59E+0−12.52E+0−0.57E+0

b7 (K mm−1 )−0.21E−2−0.34E−2−1.69E−2−5.21E−2−2.38E−2−8.78E−2−23.26E−2−8.07E−2

aO1

8.34E−3

9.08E−3

12.15E−3

15.75E−3

40.06E−3

353.72E−3

1131.76E−3

53.35E−3

aO2 (K )−0.48E−4−0.47E−4−0.61E−4−0.87E−4−2.00E−4−13.79E−4−2.26E−4−1.18E−4

aV1 (mm −1 )0.07E−3

0.18E−3

1.73E−3

5.14E−3

1.88E−3

2.91E−3

3.17E−3

8.78E−3

0.00E−5

0.00E−5−0.05E−5

0.19E−5

0.09E−5

0.24E−5

0.27E−5

0.80E−5−1−2

aV2 (mm )Table 5. RMS Error in Oxygen and Water Vapor Absorption Approximation
Freq. (GHz)

6.93

10.65

18.70

23.80

36.50

50.30

52.80

89.00

Oxygen, AO

0.00020.00020.00030.00030.00080.00620.01630.0009Vapor, AV

0.00010.00020.00110.00130.00250.00420.00460.0129

290
280
270
260
250

19 GHz

240

290

e

280
270

v

260
22 GHz

240

e

290

f

280

f

270

E

c

t

i

250

260
250

37 GHz

240
0

10

20

30

40

50

60

70

Columnar Water Vapor (mm)Fig. 4. The effective air temperature T D for downwelling radiation plotted versus the RAOB columnar water vapor. The solid
curve is the model value, and the vertical bars are the ± one standard deviation of T D derived from radiosondes.

The vapor absorption AV is primarily a linear function of V, although there is a small second order term. We
find the following expression is a good approximation for AV:
AV = aV1 V + aV2 V2(29)
Table 4 gives the aV coefficients for the 8 AMSR frequencies, and Table 5 gives the rms error in this approximation for the 8 frequencies. For the 6.9 and 10.7 AMSR channels, the rms error in this approximation is negligible,
being 0.0002 napers or less. In the 18.7 to 36.5 range, the error remains relatively small (0.001 to 0.0025 napers),
but not negligible.
The final atmospheric parameter to be specified is the vertically integrated liquid water absorption AL . Whenthe liquid water drop radius is small relative to the radiation wavelength, the absorption coefficient α L (cm−1 ) is
given by the Rayleigh scattering approximation [Goldstein, 1951]:

αL =6πρL (h )  1 − ε 
Im 

λρo
2 + ε(30)

where λ is the radiation wavelength (cm), ρL(h) is the density (g/cm3) of cloud water in the atmosphere given as afunction of h, ρo is the density of water (ρo ≈ 1 g/cm3), and ε is the complex dielectric constant of water. Notethat the dielectric constant varies with temperature and hence is also a function of h. Substituting (30) into (21)
gives

AL =

0.6πL  1 − ε Im λ2 + ε(31)

where L is the vertically integrated liquid water (mm) given byH

L = 10 ∫ dh ρL (h )

(32)

0

The leading factor of 10 converts from g/cm2 to mm. In deriving (31), we have assumed the cloud is at a constanttemperature. For the more realistic case of the temperature varying with height, ε should be evaluated at somemean effective temperature for the cloud. The specification of ε as a function of temperature and frequency isgiven in Section 2.4. An excellent approximation for (31) is found to be
A L = a L 1 1 − a L 2 ( T L − 283 ) L(33)
where TL is the mean temperature of the cloud, and the aL coefficients are given in Table 6 for the 8 AMSR frequencies. The error in this approximation is ≤ 1% over the range of TL from 273 to 288 K, which is negligible
compared to other errors such as the uncertainty in specifying the cloud temperature TL . Note that in the retrieval
algorithm, the error in specifying TL only effects the retrieved value of L. The retrieval of the other parameters
only requires the spectral ratio of AL , which is essentially independent of TL due to the fact that aL2 is spectrallyflat.
In the absence of rain, the cloud droplets are much smaller than the radiation wavelengths being considered,and equations (31) and (33) are valid. When rain is present, Mie scattering theory must be used to compute AL .
For light rain not exceeding 2 mm/h and for frequencies between 6 and 37 GHz, the Mie scattering computations
give the following approximation [Wentz and Spencer, 1998]:

Freq
(GHz)
aL1
aL2
aL3
aL4
aL5

Table 6. Coefficients for Rayleigh Absorption and Mie Scattering.
6.93
10.65
18.70
23.80
36.50
50.30
52.80
0.00780.03030.00070.00001.22160.01830.02980.00270.00601.17950.05560.02880.01130.00401.06360.08910.02810.01880.00201.02200.20270.0261
0.0425-0.00200.9546

0.36820.0236
0.0731-0.00200.89830.40210.02310.0786-0.00200.894389.00
0.96930.01460.1506-0.0020
0.7961A R = a L3 ⋅ 1 + a L4 ⋅ ( TL − 283) ⋅ H ⋅ Ra L5(34a)

The rain column height H (in km) can be approximated by:

H = 1+ 0.14 ⋅ (TS − 273) − 0.0025 ⋅ (TS − 273)2
H = 2.96

if
if

TS < 301
TS ≥ 301,

(34b)
(34c)

where TS denotes the sea surface temperature (in K). The rain rate R (in mm/h) is related to the liquid cloud water
density L by

L = 0.18 ⋅ 1 + HR .
In deriving (34a) we have used a Marshall and Palmer [1948] drop size distribution.

(34d)

2.4. Dielectric Constant of Sea-Water and the Specular Sea Surface
A key component of the sea-surface model is the dielectric constant ε of sea water. The parameter is a complex number that depends on frequency ν, water temperature TS, and water salinity s. The dielectric constant isgiven by [Debye,1929; Cole and Cole, 1941] asεS − ε∞2 j σλε = ε∞ +
−
(35)
1− η
c
1 + jλ R λwhere j = −1 , λ (cm) is the radiation wavelength, ε∞ is the dielectric constant at infinite frequency, εS is thedielectric constant for zero frequency (i.e., the static dielectric constant), and λR (cm) is the relaxation wavelength.
The spread factor η is an empirical parameter that describes the distribution of relaxation wavelengths. The lastterm accounts for the conductivity of salt water. In this term, σ (sec −1, Gaussian units) is the ionic conductivity
and c is the speed of light.
Several investigators have developed models for the dielectric constant of sea water. In the Stogryn [1971]
model the salinity dependence of εS and λR was based on the Lane and Saxton [1952] laboratory measurements ofsaline solutions. Stogryn noted that the Lane- Saxton measurements for distilled water did not agree with those ofother investigators. The Klein and Swift [1977] model is very similar to Stogryn model except that the salinity
dependence of εS was based on more recent 1.4 GHz measurements [Ho and Hall, 1973; Ho et al., 1974]. KleinSwift noted that their εS was significantly different from that derived from the Lane and Saxton measurements. It
appears that there may be a problem with Lane-Saxton measurements. However, in the Klein-Swift model, thesalinity dependence of λR was still based on the Lane-Saxton measurements. We analyzed all the measurements
used by Stogryn and Klein-Swift and concluded that the Lane-Saxton measurements of ε for both distilled waterand salt water were inconsistent with the results reported by all other investigators. Therefore, we completely exclude the Lane-Saxton measurements from our model derivation.
The model to be presented is very similar to the Klein-Swift model, with two exceptions. First, since we excluded Lane-Saxton measurements, the salinity dependence of λR is different. For cold water (0 to 10 C), our λR
is about 5% lower than the Klein-Swift value and for warm water (30 C), it is about 1% higher. Second, our valuefor ε∞ is 4.44 and the Klein-Swift value is 4.9, which was the value used by Stogryn. In the Stogryn model, η = 0,
whereas in the Klein-Swift model, η = 0.02. Grant et al. [1957] pointed out that the choice of ε∞ depends on the
choice for η, where η = 0 → ε∞ = 4.9 and η = 0.02 → ε∞ = 4.5. Thus the Klein-Swift value of ε∞ = 4.9 is probably
too high. In terms of brightness temperatures, these λR and ε∞ differences are most significant at the higher frequencies. For example, at 37 GHz and θi = 55°, the difference in specular brightness temperatures produced byour model and the Klein-Swift model differ by about ± 2 K. Analyses of SSM/I observations show that our newmodel, as compared to the Klein-Swift model, produces more consistent retrievals of ocean parameters. For example, using the Klein-Swift model resulted in an abundance of negative cloud water retrievals in cold water.
This problem no longer occurs with the new model. (The negative cloud water problem was the original motivation for doing this reanalysis of the ε model.)We first describe the die lectric constant model for distilled water, and then extend the model to the more general case of a saline solution. The static dielectric constant εS0 for distilled water has been measured by many investigators. The more recent measurements [Malmberg and Maryott, 1956; Archer and Wang, 1990] are in very
good agreement (0.2%). The Archer and Wang [1990] values for εS0, which are reported in the Handbook ofChemistry and Physics [Lide,1993], are regressed to the following expression:ε S 0 = 87. 90 exp( −0. 004585 t S )(36)
where tS is the water temperature in Celsius units. The accuracy of the regression relative to the point values for
εS0 is 0.01% over the range from 0 to 40 C.
The other three parameters for the dielectric constant of distilled water are the relaxation wavelength λR0 , thespread factor η, and ε∞. We determine these parameters by a least-squares fit of (35) to laboratory measurements
εmea of the dielectric constant for the range from 1 to 40 GHz. A literature search yielded ten papers reporting εmea
for distilled water. Values for λR0 , η, and ε∞ are found so as to minimize the following quantity:
Q = Re( ε − ε m e a )2+ I m (ε − ε m e a )

2

(37)

The relaxation wavelength is a function of temperature [Grant et al., 1957], but it is generally assumed that η andε∞ are independent of temperature. The least squares fit yields η = 0.012, ε∞= 4.44, andλ R 0 = 3.30 exp( −0. 0346 t S + 0 .00017 t 2S )(38)
These values are in good agreement with those obtained by other investigators. Our λR0 agrees with the expression derived by Stogryn [1971] to within 1%. The values for η (ε∞) reported in the literature vary from 0 to 0.02
(4 to 5). Note that using a larger value for η necessitates using a smaller value for ε∞.
The presence of salt in the water produces ionic conductivity σ and modifies εS and λR. It is generally assumed that η and ε∞ are not affected by salinity. Weyl [1964] found the following regression for the conductivityof sea water.
σ = 3.39 × 109 C 0 .892 exp (− ∆ t ζ )
(39)

(

)

ζ = 2.03× 10−2 + 1.27 × 10−4 ∆t + 2.46 ×10 −6 ∆2t − C 3.34 × 10−5 − 4.60 × 10−7 ∆t + 4.60 × 10−8 ∆2t (40)C = 0 .5536 s(41)∆ t = 25 − t S(42)
where s and C are salinity and chlorinity in units of parts/thousand. Note that we have converted the Weyl conductivity to Gaussian units of sec −1 .
To determine the effect of salinity on εS, we use low frequency (1.43 and 2.65 GHz) measurements of ε forsea water and saline solutions [Ho and Hall, 1973; Ho et al., 1974]. For the Ho-Hall data, only the real part of thedielectric constant is used in the fit. Klein and Swift reported that the measurements of the imaginary part were inerror. To determine the effect of salinity on λR, we use higher frequency (3 to 24 GHz) measurements of ε for
saline solutions [Haggis et al., 1952; Hasted and Sabeh, 1953; Hasted and Roderick, 1958]. A least-squares fit tothese data shows that the salinity dependence of εS and λR can be modeled as(

εS = εS 0 exp − 3.45 × 10 −3 s + 4.69 × 10 −6 s 2 + 1.36 × 10 −5 st S()

λR = λR 0 − 6.54 × 10 −3 1 − 3.06 ×10 −2 t S + 2.0 × 10 −4 t S2 s

)

(43)
(44)

The accuracy of the dielectric constant model is characterized in terms of its corresponding specular brightness temperature TB. For each laboratory measurement of ε, we compute the specular TB for an incidence angle of
55° using the Fresnel equation (45) below. Two T B’s are computed: one using εmea and the other using the model
ε coming from the above equations. For the low frequency Ho-Hall data, the rms difference between the ‘measurement’ TB and the ‘model’ TB is about 0.1 K for v-pol and 0.2 K for h-pol. For the higher frequency data set,
the rms difference is 0.8 K for v-pol and 0.5 K for h-pol.
Once the dielectric constant is known, the v-pol and h-pol reflectivity coefficients ρV and ρH for a specular(i.e., perfectly flat) sea surface are calculated from the well-known Fresnel equations

ρv =
ρh =

ε cos θ i − ε − sin 2 θ iε cos θ i + ε − sin 2 θ icos θ i − ε − sin 2 θ icos θ i + ε − sin 2 θ iwhere θI is the incidence angle. The power reflectivity R is then given by
R 0 p = ρp

(45a)

(45b)

2

(46)

where subscript 0 denotes that this is the specular reflectivity and subscript p denotes polarization.
An analysis using TMI data indicates small deviations from the model function for the dielectric constant of sea
water as discussed above. The effect is mainly noted in the v-pol reflectivity. In order to account for these small
differences a correction term of∆R 0 V = 4.887 ⋅ 10 −8 − 6.108 ⋅ 10 −8 ⋅ TS − 273 3
is added to the v-pol reflectivity R 0v in (46). The resulting changes in the brightness temperature range fromabout +0.14K in cold water to about –0.36K in warm water.
2.5. The Wind-Roughened Sea Surface
It is well known that the microwave emission from the ocean depends on surface roughness. A calm sea surface is characterized by a highly polarized emission. When the surface becomes rough, the emission increases

and becomes less polarized (except at incidence angles above 55º for which the vertically polarized emission decreases). There are three mechanisms that are responsible for this variation in the emissivity. First, surface waves
with wavelengths that are long compared to the radiation wavelength mix the horizontal and vertical polarization
states and change the local incidence angle. This phenomenon can be modeled as a collection of tilted facets,
each acting as an independent specular surface [Stogryn, 1967]. The second mechanism is sea foam. This mixture of air and water increases the emissivity for both polarizations. Sea foam models have been developed by
Stogryn [1972] and Smith [1988]. The third roughness effect is the diffraction of microwaves by surface waves
that are small compared to the radiation wavelength. Rice [1951] provided the basic formulation for computing
the scattering from a slightly rough surface. Wu and Fung [1972] and Wentz [1975] applied this scattering formulation to the problem of computing the emissivity of a wind-roughened sea surface.
These three effects can be parameterized in terms of the rms slope of the large-scale roughness, the fractionalfoam coverage, and the rms height of the small-scale waves. Each of these parameters depends on wind speed.
Cox and Munk [1954], Monahan and O'Muircheartaigh [1980], and Mitsuyasu and Honda [1982] derived windspeed relationships for the three parameters, respectively. These wind speed relationships in conjunction with thetilt+foam+diffraction model provide the means to compute the sea-surface emissivity. Computations of this typehave been done by Wentz [1975, 1983] and are in general agreement with microwave observations.
In addition to depending on wind speed, the large-scale rms slope and the small-scale rms height depend on
wind direction. The probability density function of the sea-surface slope is skewed in the alongwind axis and has
a larger alongwind variance than crosswind variance [Cox and Munk , 1954]. The rms height of capillary waves isvery anisotropic [Mitsuyasu and Honda, 1982]. The capillary waves traveling in the alongwind direction have agreater amplitude than those traveling in the crosswind direction. Another type of directional dependence occurs
because the foam and capillary waves are not uniformly distributed over the underlying structure of large-scale
waves. Smith's [1988] aircraft radiometer measurements show that the forward plunging side of a breaking wave
exhibits distinctly warmer microwave emissions than does the back side. In addition, the capillary waves tend tocluster on the downwind side of the larger gravity waves [Cox, 1958; Keller and Wright, 1975]. The dependenceof foam and capillary waves on the underlying structure produces an upwind-downwind asymmetry in the seasurface emissivity.
The anisotropy of capillary waves is responsible for the observed dependence of radar backscattering on winddirection [Jones et al., 1977]. The upwind radar return is considerably higher than the crosswind return. Also, themodulation of the capillary waves by the underlying gravity waves causes the upwind return to be generally
higher than the downwind return. These directional characteristics of the radar return have provided the means tosense wind direction from aircraft and satellite scatterometers [Jones et al., 1979].
To model the rough sea surface, we begin by assuming the surface can be partitioned into foam-free areas andfoam-covered areas within the radiometer footprint. The fraction of the total area that is covered by foam is denoted by f. The composite reflectivity is then given by
R = ( 1 − f ) R c l e a r + f κR c l e a r(47)
where Rclear is the reflectivity of the rough sea surface clear of foam, and the factor κ accounts for the way inwhich foam modifies the reflectivity. As discussed above, foam tends to decrease the reflectivity, and hence κ < 1.
The reflectivity of the clear, rough sea surface is modeled by the following equation:
R c l e a r = (1 − β) R g e o(48)
where Rgeo is the reflectivity given by the standard geometric optics model (see below) and the factor 1 − β accounts for the way in which diffraction modifies the geometric-optics reflectivity. Wentz [1975] showed that theinclusion of diffraction effects is a relatively small effect and hence β small compared to unity.
Combining the above two equations gives
R = (1 − F )R g e o(49)
F = f + β − fβ − f κ + fκβ(50)where F is a ‘catch-all’ term that accounts for both foam and diffraction effects. All of the terms that makeup F
are small compared to unity, and the results to be presented show that F < 10%. The reason we lump foam anddiffraction effects together is that they both are difficult to model theoretically. Hence, rather than trying to compute F theoretically, we let F be a model parameter that is derived empirically from various radiometer experiments. However, the Rgeo term is theoretically computed from the geometric optics. Thus, the F term is a measure of that portion of the wind-induced reflectivity that is not explained by the geometric optics.
The geometric optics model assumes the surface is represented by a collection of tilted facets, each acting as
an independent reflector. The distribution of facets is statistically characterized in terms of the probability density
function P(Su ,Sc) for the slope of the facets, where Su and Sc are the upwind and crosswind slopes respectively.

Given this model, the reflectivity can be computed from equation (7). To do this, the integration variables θs ,φs in(7) are transformed to the surface slope variables. The two equations governing this transformation are
k s = k i − 2(k i ⋅ n )n(51)

− S u , − Sc , 1

n =

(52)
1 + S 2u + S 2cwhere n is the unit normal vector for a given facet. Transforming (7) to the Su ,Sc integration variables yields

Rgeo∫ dS ∫ dS P(S , S ) 1 + S + S (k ⋅ n) (p ⋅ h )ρ h + (p ⋅ v )ρ v [χ(k )(1 − R ) + R ] (53)=∫ dS ∫ dS P(S , S ) 1 + S + S (k ⋅ n)
u

c

u

c

2
u

2

2
c

u

i

c

i

u

h

s

i

2
u

c

2
c

v

s

s

×

×

i

where p is the unit vector specifying the reflectivity polarization. The unit vectors hi and vi (hs and vs ) are thehorizontal and vertical polarization vectors associated with the propagation vector k i (k s) as measured in the tilted
facet reference frame. These polarization vectors in the tilted frame of reference are given by
hj =

kj ×n

(54a)

kj ×n

vj = kj ×hj(54b)

where subscript j = i or s. The terms ρv and ρh are the v-pol and h-pol Fresnel reflection coefficients given above.
The last factor in (53) accounts for multiple reflection (i.e., radiation reflecting off of one facet and then intersec ting another). χ(ks ) is the shadowing function given by Wentz [1975], and R× is the reflectivity of the secondaryintersection. The shadowing function χ(ks ) essentially equals unity except when ks approaches surface grazing
angles.
The interpretation of (53) is straightforward. The integration is over the ensemble of tilted facets having aslope probability of P(Su,Sc). The term 1 + S u2 + S c2 (k i ⋅ n ) is proportional to the solid angle subtended by
the tilted facet as seen from the observation direction specified by ki . The term (p ⋅ h i ) ρ h h s + (p ⋅ v i )ρ v v s is2

the reflectivity of the tilted facet. And, the denominator in (53) properly normalizes the integral.
To specify the slope probability we use a Gaussian distribution as suggested by Cox and Munk [1954], and weassume that the upwind and crosswind slope variances are the same. Wind direction effects are considered in Section 2.7. Then, the slope probability is given by(

P( S u , S c ) = π∆S 2) exp  − S − S 
2
u

−1



∆S 2

2
c



(55)

where ∆S2 is the total slope variance defined as the sum of the upwind and crosswind slope variances. Ocean
waves with wavelengths shorter than the radiation wavelength do not contribute to the tilting of facets and hence
should not be included in the ensemble specified by P(Su ,Sc). For this reason, the effective slope variance ∆S2increases with frequency, reaching a maximum value referred to as the optical limit. The results of Wilheit andChang [1980] and Wentz [1983] indicate that the optical limit is reached near ν = 37 GHz. Hence, for ν ≥ 37 GHz,
we use the Cox and Munk [1954] expression for optical slope variance. For lower frequencies, a reduction factor
is applied to the Cox and Munk expression. This reduction factor is based on ∆S2 values derived from the SeaSat
SMMR observations [Wentz, 1983].
∆S 2 = 5 .22 × 10 −3 Wν ≥ 37 GHz
(56a)
2
−3
1 .3
∆S = 5 .22 × 10 1 − 0 . 00748( 37 − ν ) W
ν < 37 GHz(56b)where W is the wind speed (m/s) measured 10 m above the surface. Note the Cox and Munk wind speed wasmeasured at a 12.5 m elevation. Hence, their coefficient of 5.12×10−3 is increased by 2% to account for our wind
being referenced to a 10 m elevation.
The sea-surface reflectivity Rgeo is computed for a range of winds varying from 0 to 20 m/s, for a range ofsea-surface temperatures varying from 273 to 303 K, and for a range of incidence angles varying from 49° to 57°.
These computations require the numerical evaluation of the integral in equation (53). The integration is done over
the range S2u + S 2c ≤ 4 . 5 ∆S 2 . Facets with slopes exceeding this range contribute little to the integral, and it is notclear if a Gaussian slope distribution is even applicable for such large slopes. Analysis shows that the computed
ensemble of Rgeo is well approximated by the following regression:

R geo = R 0 − [r0 + r1 θ i − 53 + r2 TS − 288 + r3 θ i − 53TS − 288]W(57)

where the first term R0 is the specular power reflectivity given by (46) and the second term is the wind-induced
component of the sea-surface reflectivity. The r coefficients are given in Table 7 for all AMSR channels. Equation(57) is valid over the incidence angle from 49° to 57°. It approximates the θi and TS variation of Rgeo with anequivalent accuracy of 0.1 K. The approximation error in the wind dependence is larger. In the geometric optics
computations, the variation of Rgeo with wind is not exactly linear. In terms of TB, the non-linear component of
Rgeo is about 0.1 K at the lower frequencies and 0.5 K at the higher frequencies. However, in view of the general
uncertainty in the geometric optics model, we will use the simple linear expression for Rgeo, and let the empirical
F term account for any residual non-linear wind variations, as is discussed in the next paragraph.
In the case of the coefficients r2 we do not use the geometric optics model coefficients (Table 7) but rather use thefollowing empirically derived forms (units are s/m-K):
r 2 V− pol = −2.1 ⋅ 10 −5(58)

r 2 h −pol = −5.5 ⋅ 10 −5 + 0.989 ⋅ 10 −6 ⋅ 37 − vifν ≤ 37(59a)r 2 h −pol = −5.5 ⋅ 10 −5

ifν > 37 .

(59b)This accounts for the observations that the wind induced emissivity is less in warm water. This effect was observed during the monsoons in the Arabian sea.

Freq. (GHz)v-pol r0h-pol r0v-pol r1

Table 7. Model Coefficients for Geometric Optics
10.65E+0
18.70E+023.80E+036.50E+050.30E+052.80E+0

89.00E+0−0.27E−03
0.54E−03−0.32E−03
0.72E−03−0.49E−03 −0.63E−03
1.13E−03
1.39E−03−1.01E−03
1.91E−03−1.20E−03
1.97E−03−1.23E−03
1.97E−03−1.53E−03
2.02E−03−0.21E−04−0.29E−04
0.44E−04−0.53E−04 −0.70E−04
0.70E−04
0.85E−04−1.05E−04
1.12E−04−1.12E−04
1.18E−04−1.13E−04
1.19E−04−1.16E−04
1.30E−04

0.75E−05

1.27E−05

1.39E−05

1.40E−05

1.15E−05−0.15E−05 −0.23E−05
0.31E−06
0.41E−06−0.36E−05
0.45E−06−0.32E−05
0.35E−06−0.30E−05
0.32E−06

6.93E+0h-pol r1
0.32E−04v-pol r2

0.11E−05

0.48E−05

0.01E−05h-pol r2
0.00E−05v-pol r3
0.00E−06h-pol r3

−0.03E−05
0.08E−06

0.00E−05−0.09E−06

0.00E−06 −0.02E−06 −0.12E−06 −0.20E−06 −0.36E−06 −0.43E−06 −0.44E−06 −0.46E−06
r0 in units of s/m, r1 in units of s/m-deg, r2 in units of s/m-K, r3 in units of s/m-deg-K

In the 10-37 GHz band, the F term is found from collocated SSM/I-buoy and TMI-buoy observations. The
procedure for finding F is essentially the same as described by Wentz [1997] for finding the wind-induced emissivity, but in this case we first remove the geometric optics contribution to R. The F term is found to be a monotonic function of wind speed described by
F = m1W
W < W1(60a)
F = m 1 W + 12 ( m 2 − m 1 )( W − W 1 ) 2 ( W 2 − W1 )W1 ≤ W ≤ W2
(60b)
F = m 2 W − 12 ( m 2 − m 1 )( W 2 + W1 )W > W2(60c)
This equation represents two linear segments connected by a quadratic spline such that the function and its first
derivative are continuous. The spline points are W1 = 3 m s and W2 = 12 m s for the v-pol and W1 = 7 m s
and W2 = 12 m s for the h-pol , respectively. The m coefficients are found so that the TB model matches the
SSM/I observations in the and TMI observations when the buoy wind is used to specify W. For the lowest channel ν = 6.9 GHz no data exist yet and we have simply used the same values as for the ν = 10.65 GHz channel.
This will be updated as soon as AMSR data become available. Table 8 summarizes the results for m 1 and m 2 at
the 8 AMSR frequencies for v and h polarizations. Both coefficients flatten out and reach a maximum forν ≥ 37 GHz.

Table 8. The coefficients m1 and m2 . Units are s/m.
Freq. (GHz) 6.93
10.65
18.70
23.80
v-pol m1
0.00020 0.00020 0.00140 0.00178
h-pol m1
0.00200 0.00200 0.00293 0.00308
v-pol m2
0.00690 0.00690 0.00736 0.00730
h-pol m2
0.00600 0.00600 0.00656 0.00660

36.50
0.00257
0.003290.00701
0.0066050.30
0.00260
0.00330
0.00700
0.00660

52.80
0.00260
0.00330
0.00700
0.00660

89.00
0.00260
0.00330
0.00700
0.00660

These results indicate that diffraction plays a significant role in modifying the sea-surface reflectivity. If diffraction
were not important, β would be 0 in equation (50), and F would be proportional to the fractional foam coverage f.
Since f is essentially zero for W < 7 m/s, m1 would be 0. This is not the case, and we interpret the m1 coefficient as anindicator of diffraction.
2.6. Atmospheric Radiation Scattered by the Sea Surface
The downwelling atmospheric radiation incident on the rough sea surface is scattered in all directions. The
scattering process is governed by the radar cross section coefficients σ o as indicated by equation (14). For a perfectly flat sea surface, the scattering process reduces to simple specular reflection, for which radiation comingfrom the zenith angle θs is reflected into zenith angle θi , where θs = θi. In this case, the reflected sky radiation is
simply RT BD . However, for a rough sea surface, the tilted surface facets reflect radiation for other parts of the sky
into the direction of zenith angle θi . Because the downwelling radiation TBD increases as the secant of the zenith
angle, the total radiation scattered from the sea surface is greater than that given by simple specular reflection.
The sea-surface reflectivity model discussed in the previous section is used to compute the scattered sky radiation
TBΩ . These computations show that TBΩ can be approximated byT B Ω = [(1 + Ω)(1 − τ )( T D − T C ) + TC ] R(61)
where R is the sea-surface reflectivity given by (49), TBD is the downwelling brightness temperature from zenith
angle θi given by (24), and Ω is the fit parameter. The second term in the brackets is the isotropic component of
the cold space radiation. This constant factor can be removed from the integral. The fit parameter for v-pol andh-pol is found to beΩ V = [ 2 .5 + 0. 018 (37 − ν) ][ ∆S 2 − 70. 0 ∆S 6 ] τ3 . 4(62a)Ω H = [ 6 .2 − 0 .001( 37 − ν )2 ][ ∆S 2 − 70 . 0 ∆S 6 ] τ 2 .0

(62b)where ν is frequency (GHz) and ∆S2 is the effective slope variance given by (56). The term ∆S 2 − 70 . 0 ∆S 6
reaches a maximum at ∆S2 = 0.069. For ∆S2 > 0.069, the term is held at its maximum value of 0.046. Ω V has alinear dependence on frequency, whereas ΩH has a quadratic dependence, reaching a maximum value at ν = 37
GHz. For ν > 37 GHz, both Ω V and Ω H are held constant at their maximum values. Approximation (62) is valid
for the range of incidence angles from 52° to 56°. For moderately high winds (12 m/s) and a moist atmosphere(high vapor and/or heavy clouds), the scattering process increases the reflected 37 GHz radiation by about 1 K forv-pol and 5 K for h-pol. At 7 GHz, the increase is much less, being about 0.2 K for v-pol and 0.8 K for h-pol.
The accuracy of the above approximation as compared to the theoretical computation is about 0.03 K and 0.2 K at
7 and 37 GHz, respectively. Note that when the atmospheric absorption becomes very large (i.e., τ is small), Ω
tends to zero because the sky radiation for a completely opaque atmosphere is isotropic.

2.7. Wind Direction Effects
The anisotropy of the sea-surface roughness produces a variation of the brightness temperature versus winddirection, as discussed in Section 2.5. In the 19 to 37 GHz band, Wentz [1992] determined this wind direction
signal using collocated SSM/I TB’s and buoy wind vectors. At an incidence angle near 53°, the wind direction
signal exhibits the following second-order harmonic variation with wind direction:∆E 1 9 − 3 7 = γ 1 cos φ + γ 2 cos 2 φ(63)
where ∆E is the change in the sea-surface emissivity and φ is the wind-direction angle relative to the azimuth-lookangle. When φ = 0° (180°), the observation is upwind (downwind). The subscript 19-37 denotes that the results
are for the 19-37 GHz band. The amplitude coefficients γ1 and γ2 are found to be essentially the same for both 19
and 37 GHz. The coefficients are different for the two polarizations and do vary with wind speed as given below
γ 1 V = 7.83 × 10 −4 W − 2 .18 × 10 −5 W 2(64a)

γ 2 V = −4 .46 × 10 −4 W + 3.00 × 10 −5 W 2(64b)

γ 1 H = 1. 20 × 10 −3 W − 8. 57 × 10 −5 W 2γ 2 H = − 8. 93 × 10 −4 W + 3. 76 × 10 −5 W 2

(65a)

(65b)
In Wentz [1992], the wind direction signal was expressed in terms of a brightness temperature change rather thanan emissivity change, and the wind speed was referenced to a 19.5 m anemometer height. In the above equations,
we have converted the Wentz [1992] expressions from ∆T B to ∆E and use a 10 m reference height for W.
Little is known about the wind direction signal for frequencies below 19 GHz. Some very preliminary datafrom the Japanese AMSR aircraft simulations suggests that the signal decreases with decreasing frequency. Other
than this, there are no experimental data on the variation of TB versus φ at 6.9 and 10.7 GHz. As an educatedguess on what will be observed at these lower frequencies we reduce the wind direction signal from its value at 19
GHz by a factor of 0.82 at 10.7 GHz and by a factor of 0.62 at 6.9 GHz.
The result for the wind direction signal from (64) and (65) should be regarded as preliminary. Recent aircraft
data Yueh et al. [1999] as well as a first analysis of TMI measurements suggest that at wind speeds below 8 m/s
the wind direction signal is noticeably smaller than the one obtained from (64) and (65), especially for the h-pol.
A reanalysis of the directional signal using data from 5 SSM/I satellites between 1987 and 1999 as well as recent
TMI data is currently under way.
3. The Ocean Retrieval Algorithm
3.1 Introduction
In general, there are three types of ocean retrieval algorithms:
1. Multiple linear regression algorithms
2. Non-linear, iterative algorithms
3. Post-launch in-situ regression algorithms
The first two types are physical algorithms in the sense that radiative transfer theory is used in their derivation.
The third type is purely statistical with little or no consideration of the underlying physics. We now describe eachof these algorithms and discuss their strengths and weaknesses.
3.2 Multiple Linear Regression Algorithm
Consider a linear process in which a set of inputs denoted by the column vector X is transformed to a set of
outputs denoted by the column vector Y. The linear process is then characterized by the matrix A that relates Y toX.
Y = AX(66)
The measurement of Y usually contains some noise ε and is denoted by~
Y = Y + e = AX + e(67)
The retrieval problem is then to estimate X given Y . The most commonly used criteria for estimating X is to find~
X such that the variance between Y and Y is minimized. Using this criteria, one finds the well known leastsquares solution:~∃ = (A TΞ − 1 A)− 1 A T Ξ − 1YX
(68)
where Ξ is the correlation matrix for the error vector ε . If the errors are uncorrelated, then Ξ is diagonal.
~
For our application, the system input vector X is the set of geophysical parameters P and the output vector Y
is the set of TB measurements. Note that X and Y can be non-linear functions of P and TB, respectively without
violating the requirement for linearity between X and Y. For example, the relationship between TB and atmospheric parameters V and L can be approximated by
TB ≈ TE 1 − R exp[− 2 sec θi ( AO + aV V + a L L )]
(69)
~

{

}

where TE is an effective temperature of the ocean-atmosphere system which is relatively constant. Then,ln( TE − TB ) = ln( RT E ) − 2 sec θi ( AO + aV V + a L L )(70)
From this we see that the relationship between TB and V, L can be linearized by transforming from Y = T B to Y =ln(TE − T B). Wilheit and Chang [1980] followed this approach and used a value of 280 K for TE . As a further
extension, Y can also include higher order terms such as TB2 and TB37V T B23H.
Likewise, the input X can be a nonlinear transformation of the geophysical parameters P. For example, thewind speed dependence of TB (i.e., ∂T B/∂W) increases with wind speed, and the relationship can be made linear
by the following transformation

W′ = W
W < W1(71a)
W ′ = W + M 1 ( W − W1 ) 2
W1 ≤ W ≤ W2(71b)
W ′ = M 2W − M 3
W > W2(71c)
This transformation represents two linear segments connected by a quadratic spline such that the function and its
first derivative are continuous.
Thus the requirement of linearity is not as constraining as it might first appear, and a generalized linear statistical regression algorithm can be represented byI
Pj = ℜ c0 j + ∑ cij ℑ(TBi ) 
i =1



(72)

where ℑ and ℜ are linearizing functions. Subscript i denotes the AMSR channel (1 = 6.9V, 2 = 6.9H, etc.), andsubscript j denotes the parameter to be retrieved (1 = TS, 2 = W, 3 = V, 4 = L). For AMSR, our initial design forthe linear regression algorithm discussed in the next section uses the following linearizing functions:ℑ( TB ) = TB
ν = 6.9 and 10.7 GHz(73a)ℑ( TB ) = − ln( 290 − TB )
ν = 18.7, 23.8, and 36.5 GHz(73b)ℜ( X ) = X(74)
After testing the initial algorithm, we will experiment with additional linearizing functions, such as the wind speed
linearization given by (71).
In principle, the cij coefficients can be found from (68) given the A matrix and the error correlation matrix Ξ .
However, even after the linearizing functions are applied, the relationship of Y versus X is not strictly linear, andthe elements of A matrix are not constant, but rather vary with P. One could find a linear approximation for the Y
versus X relationship, and then derive the cij coefficients from (68). However, we prefer the more direct approachsuggested by Wilheit and Chang [1980] in whichbrightness temperatures are computed for an ensemEnvironmental Scenes
ble of environmental scenes and then multiple linear
42,195 Radiosondes
regression is used to derive the cij coefficients, as is
5 Cloud Models
discussed in the following section.
SST Randomly Varied for 0 to 30 C
Wind Speed Randomly Varied from 0 to 20 m/s
Wind Direction Randomly Varied from 0 to 360
3.3. Derivation and Testing of the Linear Regression Algorithm
The derivation of the c ij coefficients in the
Truth: Ts, W, V, L
AMSR linear regression algorithm is shown in Figure 5. A large ensemble of ocean-atmosphere
Complete Radiative Transfer Model
scenes are first assembled. The specification of theatmospheres comes from 42,195 quality-controlled
radiosonde flights launched
Simulated AMSR TB's
from small islands during the 1987 to 1990 time
period [Wentz, 1997]. One half of these radiosonde
Gaussian Noise Added
flights are used for deriving the cij coefficients, andWithheld Data Set
the other half is withheld for testing the algorithm.
A cloud layer of various columnar water densities
Derive Coefficients for Multiple
ranging from 0 to 0.3 mm is superimposed on the
Linear Regression Algorithm
radiosonde profiles.
Underneath these simulated
atmospheres, we place a rough ocean surface. The
Algorithm Coefficientssea-surface temperature TS is randomly varied from0 to 30 C, the wind speed W is randomly varied
from 0 to 20 m/s, and the wind direction φ is ranRun Algorithm
domly varied from 0 to 360°. About 400,000 scenes
are generated in this manner.
Retrieved values for Ts, W, V, L
In nature, there is a strong correlation between
T S and W. We could have incorporated this correlaEvalulate Algorithm Peformancetion into the ensemble of the scene. For example,
we could have discarded cases of very cold waterand very high water vapor, which never occur in
Performance and Cross Talk Statistics
nature. However, for now we include these unrealistic cases in order to determine if the algorithm is Fig. 5. Derivation and testing of the linear regression algorithm

truly capable of separating the T S signal from the V signal.
Atmospheric brightness temperatures T BU and TBD and transmittance τ are computed from the radiosonde +cloud profiles of T(h), p(h), ρV(h), and ρL(h) using equations (17), (18) and (19). The reflectivity R of the roughsea surface is computed according to the equations given in Section 2.5, and the atmospheric radiation scatteredfrom the sea surface T BΩ is computed from (61). Wind direction effects are included as described in Section 2.7.
Finally, the brightness temperature T B as seen by AMSR is found by combining the atmospheric and sea-surface
components, as is expressed by (10).
Noise is added to the simulated AMSR T B’s. This noise represents the measurement error in the AMSR T B’s.
The measurement error depends on the spatial resolution. At a 60-km resolution, which is commensurate with the
6.9 GHz footprint, the measurement error is 0.1 K. A random number generator is used to produce Gaussian
noise having a standard deviation of 0.1 K. This noise is added to the simulated T B’s. At this point in the simulation, we could also add modeling error to the T B’s. Modeling error accounts for the difference between the modeland nature. It is a very difficult parameter to determine since it involves physical processes which are not sufficiently understood to be included in the current model. For now, we are not including any modeling error in thesimulations, but we will be investigating this problem in the future.
Given the noise-added simulated T B’s, the standard multiple linear regression technique is used to find the cijcoefficients. The coefficients are found such that the rms difference between Pj and the true value for the specified environmental scene is minimized. For the initial set of simulations, we use all 10 lower frequency channels(i.e., 6.9, 10.7, 18.7, 23.8 and 36.5 GHz, dual polarization). Later on, we will investigate the utility of using a
reduced set of channels.
The algorithm is tested by repeating the above process, only this time using the withheld environmental
scenes. The geophysical parameters T S, W, V, and L are computed from the noise-added T B’s using equation (72).
Statistics on the error in Pi are accumulated. The results are shown in Figure 6. The solid line in the figure shows
the mean retrieval error, and the dashed lines show the one standard deviation envelope. The retrieval error foreach of the four parameters is plotted versus the four parameters in order to show the crosstalk error matrix. The
diagonal in the crosstalk matrix verifies that the dynamic range of a given parameter is correct, and the offdiagonal plots verify that there is no crosstalk error in the retrieval.
The results look quite good. There is a little crosstalk, but it is quite small. Table 9 gives the overall rms error forthe retrievals. Wind direction variability is a major source of error in the T S retrieval. When wind direction variability is removed from the simulations, the T S retrieval error decreases to 0.3 C. The wind direction problem isfurther discussed in Sections 1.5 and 4.3.
We again emphasize that these results are very preliminary. There is much more work to do. For example,
the cloud models need to be more variable and the performance of the relatively simple LSR algorithm needs tobe compared with the non-linear algorithm discussed in the next section.
Table 9. Preliminary Estimate of Retrieval Error
Ocean Parameter
Rms Error
Sea-Surface Temperature
0.58 CWind Speed
0.86 m/s
Columnar Water Vapor
0.57 mm
Columnar Cloud Water
0.017 mm
3.4. Non-Linear, Iterative Algorithm
The major shortcoming of the multiple linear regression algorithm is that the non-linearities in the T B versus P
relationship are handled in an ad hoc manner. The linearization functions are only approximations, and the inclusion of second order terms such as T B2 and T B37V T B23H do not really describe the inverse of the T B versus P relationship. A more rigorous treatment of the non-linearity problem is to express the TB versus P relationship interms of a non-linear model function F(P), and then invert the following set of equations
TBi = Fi (P ) + ε i(75)
where subscript i denotes the observation number and ε i is the measurement noise. The number of observations
must be equal to or greater than the number of unknown parameters (i.e., the number of elements in P). For each
set of AMSR observations, equations (75) are inverted to yield P. This method is much more numerically intensive than the linear regression algorithm in which there is a fixed relationship between P and T B. However, given
today’s computers, the computational burden is no longer a problem.

Equation (75) is solved using an extension of Newton’s iterative method. In Newton’s method, the model function is expressed as a Taylor expansion:

T Bi = Fi ( P) + ∑ (Pj − Pj )4

j =1

SST DIF (C)WIND DIF (M/S)

VAPOR DIF (MM)

CLOUD DIF (MM)

0

30 0

∂Fi
∂Pj

+ O 2 + εi

(76)

P

20 0

70 0

.3

.1

.1

0

0

-.1

-.1

3

3

0

0

-3

-3

2

2

0

0

-2

-2

2

2

0

0

-2

-2
0

30 0
SST (C)

20 0
WIND (M/S)

70 0
VAPOR (MM).3
CLOUD (MM)Fig 6. Preliminary results for the linear statistical regression algorithm for AMSR. The solid line in the figure shows the mean retrieval error, and the dashed lines show the one standard deviation envelope. The retrieval error for each of the four parameters isplotted versus the four parameters in order to show the crosstalk error matrix. The diagonal in the crosstalk matrix verifies that the
dynamic range of a given parameter is correct, and the off-diagonal plots verifies that there is no crosstalk error in the retrieval.

where P is a first guess value for P and O2 represents the higher order terms in the expansion. This system ofsimultaneous equations is written in matrix form as∆TB = A ∆P + O 2 + e(77)
where A is a matrix of i × j elements and ∆ TB , ∆ P, and ε are column vectors. The elements of A, ∆ TB, and ∆ P
are

Aij =

∂Fi
∂Pj

(78)
P

∆TB i = TB i − Fi ( P )∆Pj = Pj − Pj

(79)
(80)

Equation (77) is solved by ignoring the higher order terms (i.e., by setting O2 to zero), and the solution is

P = P + (A TΞ − 1 A) − 1 ATΞ − 1 ∆ TB

(81)

where Ξ is the error correlation matrix. This procedure is then repeated with P from (81) replacing P , and several
such iterations are performed. For the no-noise case ( ε = 0), Ξ drops out of the formulation and an exact solution
is obtained when ∆ TB goes to zero. For the noise case, a solution is found when ∆ TB reaches some constant
minimum value.
The solution of P can be constrained by the inclusion of a priori information. This is accomplished by including additional equations in (77). For example, if ancillary information on wind direction were available, then thefollowing equation could be added to (77)φ = φ∃ + ε φ(82)
where φ∃ is the a priori estimate of φ and ε φ is the rms uncertainty in that estimate. Similar constraining equations
can be included for other types of information such as the vertical distribution of water vapor and air temperature.
In general, there is no guarantee that a solution will be found using this method. Furthermore, if a solution isfound, there is no guarantee that it is an unique solution. However for the case of AMSR, the relationships between P and TB are quasi-linear in that ∂T B/∂P > 0 for all channels except 36.5 GHz in cold water, for which∂T B/∂T S is < 0. Experience has shown that a solution is nearly always found. It also appears that this solution is
unique, but this needs to be verified.
We have been assuming that the TB versus P relationship can be exactly described by a non-linear model
function F. In this case, the non-linear, iterative algorithm has the distinct advantage of finding the exact solution.
In comparison, the P found by the linear regression algorithm would be in error by some degree due to the nonlinearities. However, in practice it is not possible to exactly represent the TB versus P relationship in terms amodel function F(P). For example, the TB not only depends on the columnar content of water vapor but also on
vertical distribution of the vapor. Thus, some approximations need to be made when going from the integral
equations of radiative transfer to a simplified model function F(P). These assumptions were discussed in length in
Section 2. In the derivation of the linear regression algorithm, the complete integral formulation of the radiativetransfer theory is used, and there is no need for the simplifying assumptions.
In comparing the two types of algorithms, there is a tradeoff between errors due to non-linearities in the linear regression algorithm and errors due to simplifying assumptions in the non-linear, iterative algorithm. Our plan
is to develop and test both types of algorithms in parallel, compare their relative performance, and then select one
or the other.
3.5. Post-Launch In-Situ Regression Algorithm
After AMSR is launched, purely statistical algorithms can be developed by collocating the AMSR TB’s withselected in-situ sites. A simple least-squares regression is then found that relates the in situ parameter to the T B’s.
The mathematical form of this type of algorithm is identical to the linear regression algorithm given by (72). The
difference is that the cij coefficients are not derived from radiative transfer theory, but rather from the regression to
the in situ data. Examples of this type of algorithm are the Goodberlet et al. [1989] SSM/I wind algorithm and theAlishouse et al. [1990] SSM/I water vapor algorithm.
The strength of the purely statistical algorithm is that it does not require a radiative transfer model, and hence
it is not affected by modeling errors. The weaknesses are:
1. The algorithm for AMSR cannot be developed until after launch.
2. Large in situ data sets covering the full range of global conditions must be assembled and collocated with the
AMSR observations.

3. The purely statistical algorithm is keyed to specific sensor parameters such as frequency and incidence angle.
For example, none of the algorithms developed for SSM/I can be applied to AMSR. In contrast, SSM/I algorithms based on radiative transfer theory can be interpolated to the new AMSR frequencies and incidence angle.
4. Cross-talk among the various geophysical parameters is a problem for the statistical algorithm. For example,
consider sea-surface temperature TS and water vapor V which are highly correlated on a global scale. A purely
statistical algorithm will mimic this correlation and will generate a TS product that is always highly correlatedwith V. In nature, when the true V changes and T S remains constant (i.e., a weather system passing by), the statistical algorithm will erroneously report a change in TS.
5. For the case of cloud water retrieval, for which there is no reliable in situ data, this type of algorithm cannot beused.
We think it is a mistake to ignore the physics when developing an algorithm. It may be the case that relatively simple regressions can be used to retrieve some of the parameters. However, it is important that these regressions be understood in the context of radiative transfer theory. Thus, after AMSR is launched and the collocated in situ data are available, we will calibrate the pre-launch algorithm by making small adjustments to the radiative transfer model rather than developing purely statistical algorithms. This calibration activity is discussed in
the Section 5.
3.6. Incidence Angle Variations
The retrieval of sea-surface temperature and wind speed are sensitive to incidence angle variations. A 1° error in specifying θi produces a 6 C error in TS and a 5 m/s error in W. Thus, it is crucial that the incidence angle
be accurately known and that the retrieval algorithm accounts for incidence angle variations.
The pointing knowledge for the PM platform is specified to be 0.03°/axis. This figure is the 3-standard deviation error in the know ledge of the roll, pitch, and yaw. Yaw variations do not affect the incidence angle, but rolland pitch do. The corresponding 3-standard deviation error in incidence angle is approximately 0.05°. The retrieval accuracy for the geophysical parameters are in terms of a 1-standard deviation error, so we convert the incidence angle error to a 1-standard deviation error of 0.016°, and this results in a 0.1 C error in the TS retrieval anda 0.1 m/s error in the W retrieval. The specification of pointing knowledge for the PM platform is, therefore, sufficient. However, the pointing knowledge of the AMSR instrument is yet to be specified. We will be paying
close attention to this instrument specific ation.
In the non-linear, iterative algorithm, incidence angle is an explicit parameter in the model function, andhence θi variations are easily handled by simply assigning a value to θi before doing the inversion process. There
are two possible methods for handling incidence angle variation in the linear regression algorithm. First, include
incidence angle as an additional term in the regression or second, normalize the TB’s to some constant incidence
angle, say 55°, before applying the regression. This normalization is expressed by
TB 55° = TB (θ i ) + µ ⋅ θ i − 55°(83)
where µ represents the derivative ∂T B/∂θi , which depends on the TS, W, V, and L. We find that µ can be accurately approximated by a TB regression of the type given by (73). This method works well when the incidence
angle variations are ± 1° or less, which should be the case for AMSR.

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