Occurs when surface water becomes more dense and sinks to the bottom of the lake. Is driven by wind, the Coriolis effect, and Ekman transport. Where Ekman transport moves surface waters toward the coast, the water piles up and sinks in the process known as downwelling. This example is from the.
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Bradley, in, 2015 Short-Wave RadiationDownwelling short-wave radiation at the surface has a component due to the direct solar beam, and a diffuse component scattered from atmospheric constituents and reflected from clouds. Upwelling short-wave radiation comes from reflection at the surface. Both can be measured with the pyranometer, facing either upward or downward (F and G in Figure 1); their ratio is the surface albedo. The pyranometer sensor is a blackened horizontal surface on which the radiation falls, bonded to a thermopile whose reference junction is the instrument body.
Accuracy for the instrument is usually quoted as 2%. The direct solar beam can be measured with a pyrrheliometer. These are more accurate, but are generally unsuitable for operational use because they need to track the sun. Their main use is as calibration standards for working instruments.The diffuse component is obtained with a ‘shadow-band’, set to shield the sensor from the direct solar beam. The position of the band is changed manually to follow the annual variation of solar elevation.
Instruments have also been built with a rotating shadow band which alternately shields and exposes a fast-response radiation sensor, such as a solid state photocell. This system is less accurate, but useful on ships or other moving platforms. Helen Cleugh, Sue Grimmond, in, 2012 3.2.3.1 Urban Radiation BalanceThe downwelling shortwave radiation (diffuse plus direct beam radiation) is reduced in urban areas, varying from minor reductions (∼5%) in those cities with low aerosol concentrations to much larger reductions (up to ∼30%) in cities that have high levels of particulate pollution, for example, Hong Kong (Stanhill and Kalma, 1995) and Mexico City (Jauregui and Luyando, 1998); see also references in Stanhill and Cohen (2009).
Air pollutants in the urban atmosphere alter the transmission of solar radiation, causing increased diffuse radiation and attenuation of both UV-B and photosynthetically active radiation (Arnfield, 2003).With increased emissivity due to the presence of GHG (especially CO 2), aerosols, elevated temperature, and often humidity levels (see discussion below on the UBL), the urban atmosphere has the potential to create an enhanced greenhouse effect, very similar to that occurring at the global-scale, which increases the flux of downwelling longwave. As for solar radiation, these changes in longwave radiation vary from city to city.The net result of lower downwelling shortwave, and increased longwave, radiation fluxes is little change in the total incoming all-wave radiation (Oke, 1988; Cleugh, 1995a; Arnfield, 2003). Any reduction in solar radiation receipt is very likely to be offset by the lower albedo of the urban canopy, which has been well-established through many measurement programmes to be between 0.10 and 0.15 for suburban and urban landscapes (Cleugh, 1995a; Jin et al., 2005). These effects, and the larger active area for radiation absorption, means that urban areas absorb more shortwave radiation than a non-urban landscape subjected to the same global radiation. Similarly, the reduced SVF in urban canyons impedes longwave radiation losses and so the receipt of net longwave radiation is typically increased.The net effect of these changes in the component radiation fluxes is that differences in the net all-wave radiation flux between urban and plant canopies are, in general, surprisingly small (Cleugh, 1995a; Arnfield, 2003; Lietzke and Vogt, 2009; and examples for particular cities are included in Christen and Vogt, 2004; Rotach et al., 2005; Loridan and Grimmond, 2011). (2.73) E d ( λ, 0 − ) = ( 1 − ρ dd ) E dd ( λ ) + ( 1 − ρ ds ) E ds ( λ ) + ρ u E u ( λ, 0 − ).Not all incident radiation penetrates into the water, but the surface reflects a portion back to the upper hemisphere. The radiance reflection factors given by the Fresnel equation Eq.
(2.70) average to irradiance reflection factors depending on the incident radiance distribution, slope distribution of the water surface facets, and solar elevation. For clear sky conditions they are in the order of ρ dd≈0.02–0.03 for the direct component ( Jerlov, 1976; Preisendorfer and Mobley, 1986), ρ ds ≈ 0.06–0.07 for the diffuse component ( Gregg and Carder, 1990), and ρ u ≈ 0.50–0.57 for the upwelling irradiance ( Jerome et al., 1990; Mobley, 1999). Illustration of variability of downwelling irradiance in water. The measurements were made within 100 s just below the water surface. Adopted from Gege and Pinnel (2011).Since E d is used to derive AOPs like K d, R, or r rs, the wave focusing effect can introduce large errors to underwater measurements of these parameters.
![Upwelling Upwelling](http://oceanservice.noaa.gov/facts/upwelling.jpg)
The errors can be reduced by averaging sufficient measurements, or by applying a spectral model of the induced E d changes for correction ( Gege and Pinnel, 2011; Gege, 2012a). Geometry based models are useful for simulation, but inapt for data analysis since the actual water surface geometry cannot be determined with sufficient accuracy.
The downwelling and upwelling irradiances (radiant flux per unit of area, see Radiative Transfer in the Ocean) are convenient measurements to make at sea to characterize the penetration of daylight into the water column. The depth variations of these quantities are quantified by diffuse attenuation coefficients ( Radiative Transfer in the Ocean).
When dealing with downward irradiance E d, the corresponding coefficient is K d (simply written K). It depends on the IOPs of the water and on the geometrical structure of the light field. Only by approximation, it can be seen as the sum of a term due to the water itself and a varying contribution of all materials (particulate and dissolved) originating from biological activity, so that. On the basis of many field measurements in oceanic waters, it has been shown that the spectral K bio( λ) values do not vary at random but are interrelated. The proposed optical classification is based on the realization that such a rather regular change affects simultaneously all the wavelengths and progressively modifies the entire spectrum. Bio-optical algorithms derived from statistical analyses of these field data allow the entire K( λ) spectrum to be specified, as soon as K( λ 0), at a reference wavelength λ 0 is known ( Table 3, eqn 3.1).
Diffuse attenuation coefficient for downward irradiance, at selected wavelengths, as a function of the chlorophyll concentration in Case 1 waters. The initial values (the ordinates when Chl=0.02 mg m −3) are almost the pure sea water values ( K w in eqn 7, see also eqn 3.1, in Table 3, for the term K bio).To the extent that K is largely determined by absorption, the nonlinear character of the correlation between K and Chl is not surprising and resembles that observed for a ( Figure 3B). By integrating over the whole visible domain (the photosynthetic available radiation (PAR) domain, see Radiative Transfer in the Ocean), a relationship between Z eu, the depth of the euphotic zone (cf. Table 1), and Chl can be obtained. This nonlinear relationship can also be derived through a direct analysis of the column-integrated chlorophyll content and of Z eu, observed at sea by using a photometer able to determine the vertical PAR profile. This bio-optical algorithm is useful to predict, in Case 1 waters, the depth of the euphotic zone when the vertical chlorophyll profile has been determined (eqn 3.3, Table 3).
(4.2.21) LW ↓ C = LW ↓ ( 1 + 0.1762 C 2 )fits their data well with a near zero bias and a rms of 14 W m −2. None of the other formulas that they used for calculating the LW flux had as small a bias and rms error.Josey et al., (1997) have compared measurements of downwelling LW flux in the North Atlantic and the Southern Ocean to estimations based on Clark et al., (1974), Bunker (1976), and Bignami et al., (1995) formulas, and conclude that the Clark et al., (1974) formulation for the net LW flux led to the least bias and recommend its use in midlatitudes. It contains a latitude-dependent cloud cover coefficient. 4.13 t plume, j, g = e − c g h cosθ k g ≈ 1 − c g h μ k gwith c g (ppmv) the volume mixing ratio of the gas (ratio between the volume of gas and the volume of the air), integrated into the layer with a thickness h (the length of the path l = h/cos (θ)), k g the spectral absorption coefficient of the gas being studied (m - 1 ppmv - 1), θ the view zenith angle for the upwelling transmission and the solar zenith angle for the downwelling transmission and μ the cosine of the zenith angle being considered.
Figure 4.10 shows, for different gases of interest having non-null absorption in the reflective domain, the value of the upwelling transmission corresponding to an integrated value of 1,000 ppm.m, whether 1 ppm of 1,000 m thickness or 10 ppm of 100 m thickness. We can see that detectability is a function of the absorptivity of the gas.
It should be noted that the case of CO 2 is distinctive, since it is naturally present in relatively strong concentrations in the atmosphere, typically 400 ppm throughout the troposphere. Zhang, in, 2015 WavesThe strong surface westerly wind of the MJO often generates downwelling oceanic Kelvin waves that propagate eastward and equatorial Rossby waves that propagate westward along the equatorial waveguide. The downwelling Kelvin waves play an instrumental role in MJO effects on ENSO, IOD, and ITF, among others. When these Kelvin waves reach the eastern boundary of an ocean basin (e.g., Sumatra), they lead to coastal waves that move along the coastlines toward higher latitudes where they cause local impacts, such as large intraseasonal fluctuations in SST near the northwestern coast of Australia.
Swift, in, 2011 11.2.1 Mean Wind ForcingThe mean winds in the Southern Hemisphere result in Ekman downwelling over the broad latitude region from 50°S to 10°S ( Figure 5.16d and Figure S11.3a on the textbook Web site). This produces Sverdrup forcing for a standard, anticyclonic subtropical gyre ( Figure 5.17 and Figure S11.3b). The gyre forcing is different from that for the South Pacific and South Atlantic, because the southern cape of Africa, at about 35°S, lies well within the major subtropical gyre forcing. The subtropical gyre “runs out” of western boundary before it “runs out” of wind forcing for the gyre. Consequently, the western boundary current, the Agulhas, overshoots the tip of Africa, making it different from the other four subtropical gyre western boundary currents. The wind forcing then continues the subtropical circulation far to the west to the coast of South America, where there is a southward western boundary current (Brazil Current).
The actual circulation is much more complex as the Agulhas turns back to the east after it separates from the African coast, shedding large eddies at the retroflection that propagate westward into the South Atlantic rather than continuing westward as a smooth flow to the coast of South America. In any case, the wind forcing ensures that the subtropical gyres of the Indian and South Atlantic are connected.At the eastern side of the Indian Ocean's subtropical gyre region, there is some connection with the South Pacific's subtropical circulation. East of Tasmania, the subtropical circulation is more part of the South Pacific's circulation, although part of the East Australian Current (EAC) leaks into the Indian Ocean circulation (Section 10.4.1).The mean winds in the tropical and northern Indian Ocean produce a net upwelling region between the equator and 15–20°S. This is associated with a cyclonic gyre consisting of the westward SEC on the south side, the eastward South Equatorial Countercurrent (SECC) on the north side, and a northward western boundary current (East African Coastal Current; EACC).The Southwest Monsoon, producing net downwelling and Sverdrup transport forcing for a mean anticyclonic circulation, dominates in the mean winds in the Arabian Sea. The Northeast Monsoon regime, though, is quite different (see next section). (5) ε = 0.1949 + 0.107510 ε 10 + 0.066411 ε 11 + 0.123312 ε 12 + 0.392513 ε 13 + 0.1111 ε 14where ε 10, ε 11, ε 12, ε 13, and ε 14 are the five ASTER channels’ emissivities.Accuracy of the ground LST estimated by Eq.
(4) depends on the accuracy of the upwelling and downwelling radiation and the broadband emissivity. For the field pyrgeometers, both for upwelling and for downwelling measurements, the uncertainty is estimated to be ± 5 W m − 2 ( Augustine and Dutton, 2013). The resulting overall LST uncertainty lies in a range between 0.6 and 2 K for all stations, which translates to a relative uncertainty smaller than 1% of the LST value (Glob Temp Validation Report).In order to be globally representative, stations over several major surface types from radiation networks of SURFRAD, BSRN, and GMD were used.
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