Albedo and Reflectivity

Albedo

Albedo is the percentage of solar radiation reflected by an object. The term is derived from the Latin albus, white. A pure white object would reflect all radiation that impinges on it and have an albedo of 100%. A pure black object would absorb all radiation and have an albedo of 0%. Bright Earth features such as clouds, fresh snow, and ice have albedos that range from 50% to 95%. Forests, fresh asphalt, and dark soils have albedos between 5% and 20%. Table A15 presents representative albedos for a variety of objects. Knowledge of albedo is important because absorbed solar radiation increases the amount of energy available to the Earth’s surface and atmosphere, whereas reflected radiation returns to space.

Appreciation of the relation between albedo and climate extends historically to at least classical Greek times. P. Bouguer and J. Lambert first formulated the principles and theories by which albedo and reflectivity may be explained and measured in the eighteenth century, but accurate measurements did not begin until the early twentieth century (Fritz and Rigby, 1957). The work of early investigators, including A. Ångström, C. Dorno, N. Katlin, F. Götz, H. Kimball, and others, rapidly developed an extensive body of knowledge concerning albedos that is still drawn on today (see annotated bibliography by Fritz and Rigby, 1957). One of the more interesting approaches to early observations of the Earth’s planetary albedo employed measurements of earthshine and sunshine on the moon (Danjon, 1936, cited in Fritz and Rigby, 1957). Similar efforts continue today (Goode et al., 2001). Use of aircraft and spacecraft as observing platforms has significantly expanded albedo studies in recent decades (Barrett, 1974; Brest and Goward, 1987; Schaaf et al., 2002).

Table A15 Albedos for selected objects

Reflectivity

Reflectivity is the capacity of an object to reflect solar radiation. It is described as a function of radiation wavelength and is determined by the physical composition of the object. The adjective “spectral” is frequently used in conjunction with reflectivity to indicate that reflectivity varies as a function of solar wavelength.

Representative spectral reflectivity measurements for common Earth surface features are given in Figure A22. Soil reflectivity in general increases monotonically with increasing wavelength to about 1.3 µm and then decreases with sharp dips at 1.4 µm and 1.9 µm because of absorption by soil water. Living green vegetation reflectivity is low in the visible portion of the spectrum (0.4–0.7 µm) as a result of absorption by chlorophyll and related pigments, high in the near infrared (0.7–1.3 µm) because of light scattering by internal leaf cellular structures, and decreases past 1.3 µm in a manner similar to soils due to absorption by water within leaves. Snow is highly reflective in the visible and decreases to low values in the infrared, again as a result of water absorption. Water reflects little radiation in any portion of the spectrum when solar elevation is high.

Figure A22

Spectral reflectance of selected Earth surface features.

Spectral reflectivity varies significantly for each surface type as a function of physical condition and composition. Soil reflectivity varies because of variations in moisture content, particle size, organic matter content, surface roughness and mineral composition. Figure A23 presents the variation of a silty loam soil reflectivity due to changes in moisture content. Vegetation reflectivity varies with percentage ground cover, canopy geometry, leaf size, and area and plant growth stage. Snow reflectivity varies with crystal size, compaction, age, and liquid water content. Water reflectivity is affected by turbidity, depth, and phytoplankton concentrations. Also, because water in its pure form is a dielectric, its albedo increases as the angle of incidence of radiation decreases. Water albedo is lowest when the sun is near zenith and increases to near 100% when the sun is near the horizon. (For further discussion see chapter 4, Kondrotyev, 1973; chapter 8, Miller, 1981). Other factors may affect the reflectivity of these surfaces, such as lichen crusts, and other materials such as rocks, and man-made materials (e.g. asphalt and concrete) also display unique spectral reflectivity patterns.

Figure A23

Variations of silty loam soil spectral reflectance as a function of water content (percentage water content shown by each plot) (after Kondrotyev, 1973).

Relation of reflectivity to albedo

Albedo is the integrated product of incident solar radiation spectral composition and the spectral reflectivity of the object. Outside the atmosphere, solar radiation spectral composition is relatively constant, peaking at about 0.5 µm, decreasing rapidly at shorter wavelengths to small amounts at 0.2µm, and decreasing less rapidly at longer wavelengths to small amounts at about 4.0 µm.

The atmosphere selectively absorbs and scatters solar radiation. As a result, at the Earth’s surface the spectral composition of solar radiation varies significantly as a function of atmospheric conditions (e.g. clouds, water vapor, and dust) and solar elevation (Robinson, 1966; Dickinson, 1983). The majority of albedo measurements have been carried out under clear-sky, high-sun elevation conditions (Table A15). Under cloudy conditions radiation is predominantly in the visible spectrum. This decreases the albedos of soils and vegetation but increases snow albedo (Miller, 1981). When atmospheric turbidity is high, or the sun is low in the sky, the spectral distribution of solar radiation shifts to the red and infrared portion of the spectrum. Soil and vegetation albedos increase and snow albedo decreases (Kondrotyev, 1973). This variability points out the need to know both the spectral reflectivity of objects and the spectral composition of incident radiation in order to evaluate earth albedos.

Surface and planetary albedos

Two global albedo measurements are of general interest to climatologists: surface and planetary albedos. Surface albedo is the ratio of incident to reflected radiation at the interface between the atmosphere and the Earth’s land and water areas. Almost 75% of all solar energy absorbed by the Earth is absorbed at this interface; the remainder is absorbed in the atmosphere (Sellers, 1965). Any change in surface albedo will alter climate by significantly changing the amount of solar energy absorbed by the planet. Several studies have evaluated the Earth’s surface albedo with resultant estimates ranging between 13% and 17% (Sellers, 1965; Kondrotyev, 1973; Hummel and Reck, 1979; Briegleb and Ramanathan, 1982). This range of results is suggestive of current limitations in knowledge of the distribution and reflectivity of Earth surface features.

Planetary albedo is the ratio between incident and reflected radiation at the top of the atmosphere, also referred to as “Bond” albedo, named after the astronomer who first described this metric. It includes the effects of reflection from the atmosphere, particularly clouds, and surface albedo. Only about 6% of the incident radiation is reflected by scattering in the atmosphere but, on average, 24% of incident radiation is reflected by clouds (Sellers, 1965). Changes in atmospheric turbidity or cloud cover can alter climate by changing the amount of solar radiation that reaches the Earth’s surface. Studies of the Earth’s planetary albedo have been carried out over the past 60 years (Barrett, 1974). Barrett notes that estimates have progressively decreased from 50% in early studies to current estimates between 30% and 35%, based on satellite estimates. He suggests that this trend is due to improved knowledge of global cloud cover. However, the possibility of inter-annual and longer-term variations in planetary albedo should not be overlooked (Rossow and Zhang, 1995).

Geographic patterns

Both surface and planetary albedos increase with distance from the equator (Figure A24). Surface albedo shows a slight minimum at the equator because of dense evergreen forest and a secondary minimum at 40°S latitude because of the large extent of open ocean at this latitude. Increases at 25° to 30° north and south result from the presence of subtropical deserts with relatively high albedos (>40%). At higher latitudes the seasonal or permanent occurrence of snow cover and sea ice raise average albedos, which are in excess of 60% in the polar regions. The latitudinal patterns of planetary albedo are less extreme than the surface trends. Two factors affect this difference. Tropical and midlatitude land areas that are vegetated are in latitudes of frequent cloud occurrence. Contrasts between vegetated and desert latitudes are thus less apparent in the planetary figures. In addition, interactions between surface albedo and atmospheric scattering and absorption tend to reduce albedo differences between the tropics and the poles. Where surface albedo is high, the reflected radiation passes back and forth through the atmosphere many times, increasing absorption both at the surface and in the atmosphere. Over regions of low albedo, scattering in the atmosphere increases planetary reflectance when compared to surface albedos (Hummel and Reck, 1979).

Figure A24

Latitudinal variations in surface and planetary albedo (after Hummel and Reck, 1979).

Measurements

Traditionally, two pyranometers, one pointed toward the sky, the other toward the surface, have been used to measure albedos. Pyranometers are solar radiation measurement devices that respond thermally to record the amount of radiation incident on the device. They are designed to absorb all wavelengths of solar radiation equally (Sellers, 1965). In the late 1960s, photoelectric detectors, such as silicon cells, became widely used for albedo and spectral reflectance measurements (Dirnhirm, 1968). The advantage of photon detectors is that they respond quickly to changes in incident radiation and thus permit high-resolution measurements, particularly from aircraft and spacecraft (Barrett, 1974; Justice and Townshend, 2002). One limitation of photon detectors is that they are sensitive to restricted spectral ranges. Silicon, for example, senses only wavelengths between 0.5 µm and 1.0 µm. Measurements must be either compensated for in those portions of the solar spectrum not observed, or two or more different detectors must be used. However, ease of use in the field and in aircraft and spacecraft has significantly increased their use for albedo measurements.

Human effects on Earth’s albedo

Recently investigators have suggested that human modifications of the Earth’s surface, accompanying continued expansion of urbanization, agriculture and forestry, may be altering the planet’s albedo. For example, Otterman (1977) showed that overgrazing in desert regions can increase surface albedo by as much as 20%. Charney (1975) estimated that such changes may suppress rainfall, which would enhance the process of “desertification” that is occurring in sub-Saharan Africa. Sagan et al. (1979) proposed that extensive deforestation in tropical rainforests may significantly increase surface albedo and result in major climate changes. Such change may influence local climatic conditions, but global assessment suggests that such human changes to the Earth’s land area contribute only slightly to possible global albedo changes (Henderson-Sellers and Gornitz, 1984).