Snow albedo

The influence of snow and ice in the climate system stems largely from their high albedo. Shortwave radiation is backscattered from snow and ice, with the reflectivity strongly dependent on wavelength. Neglecting this complexity for now, the broadband surface albedo can be defined from the ratio of reflected to incoming shortwave radiation, a = QS/QS.

A range of albedo values for snow and ice is given in table 2.1. Albedo is highly variable, spatially and temporally, so it is not recommended to adopt a single, constant value for modeling of snow and ice in the climate system. As an illustration, figure 2.4 plots the evolution of surface albedo through the melt season as measured on Haig Glacier. Fresh snow has an albedo of 0.8-0.9. This is typical of seasonal snowpacks during winter months and year-round values in the accumulation area of the polar ice sheets. Snow albedo typically decreases in old snow-packs, particularly when temperatures are above 0°C and meltwater is introduced into the snow. Albedo values of a mature, wet snowpack are closer to 0.6 and can fall well below this (figure 2.4).

Calendar day (2004)

Figure 2.4. Measured surface albedo evolution through the melt season at Haig Glacier in the Canadian Rockies. The transition from seasonal snow cover to bare glacial ice is evident from the albedo drop in early August. High-albedo spikes in the record are associated with snowfall events.

Calendar day (2004)

Figure 2.4. Measured surface albedo evolution through the melt season at Haig Glacier in the Canadian Rockies. The transition from seasonal snow cover to bare glacial ice is evident from the albedo drop in early August. High-albedo spikes in the record are associated with snowfall events.

Several processes are responsible. Snow-grain meta-morphism generally causes an increase in crystal size in the days and weeks after a fresh snowfall. The effective optical radius of spherical snow particles is ca. 50 mm for fresh snow, increasing to 100 mm within days and 1 mm or more in mature, melting snowpacks. This increase in grain size increases the path length for solar radiation transmit-tance in the near-surface snow layer, effectively reducing the incidence of scattering reactions at intergranular snow-air interfaces and reducing the snow albedo.

There are compounding effects beyond just grain meta-morphism in melting snowpacks. Chemical impurities and meltwater content also reduce snow albedo. It requires only a few parts per million of impurities to cause an albedo reduction of several percent. The effect depends on the type of impurity, with black carbon (soot) inducing a discernible impact for concentrations of order 0.1 ppm, comparable with mineral dust levels of 10 ppm. Natural snowpacks in tropical and midlatitude areas have impurity concentrations of this order of magnitude or greater, but dust concentrations in polar snowpacks are commonly less than 0.1 ppm.

The effects of grain-scale recrystallization may saturate after several days or weeks, but meltwater effects and impurity concentration increase throughout the melt season. The result is a reduction in albedo to late-summer values as low as 0.3 in midlatitude snowpacks. Values vary between sites and from place to place in a given snowpack, as a function of the concentration and type of impurity. The role of liquid water is unclear in scattering reactions; because the refractive index of snow crystals and liquid water is similar, there should be little effect. However, water content may increase the effective radius of grains. On a macroscale, ponded surface water causes significant albedo reductions on snow and ice surfaces, accelerating spring and summer melt on lake ice, sea ice, and glaciers.

Ice typically has a lower albedo than snow, but estimates in the literature again vary substantially, from 0.1 to 0.6. A value ai = 0.2 is typical of midlatitude glaciers, whereas ai = 0.5 is more representative of sea ice and the ablation zones of polar glaciers and ice sheets. Crystal size, impurity concentration on the glacier or sea-ice surface, liquid water, and superimposed ice content again play large roles, which gives rise to a large amount of spatial variability; micro-topography causes some areas to pond water and debris, whereas other areas are flushed. On glacier surfaces, albedo generally decreases at lower elevations in the ablation zone due to higher melt rates and longer exposure times for the surface. These influences contribute to higher debris concentrations within a given melt season and cumulatively, over many years. The albedo of debris-rich ice is 0.1-0.2. Clean ice has values closer to 0.4. Superimposed (refrozen) ice and blue ice are much brighter, with values closer to 0.6. Blue ice is found in the ablation area of polar ice caps, where wind scour can create bare-ice zones.

Albedo values for lake and sea ice depend on the age, type, and thickness of the ice. Young ice that is only a few centimeters thick is highly transparent to solar radiation, so absorption in the underlying water reduces the effective spectral albedo to values of 0.1-0.2. Young "gray ice" that is up to 30 cm thick is still relatively dark, with reported albedo values of 0.3-0.4. Ice becomes opaque as it thickens beyond this, and there is also more complete ice cover in an area (versus a large fraction of open water for young, growing ice), so albedo values of developed first-year or multiyear ice are in the range 0.6-0.7, dropping during the melt season in the presence of melt ponds. When sea ice has a snow cover exceeding a few centimeters, snow spectral properties dominate the albedo, so values of 0.8-0.9 are typical.

chapter 2

In addition to these spatial and temporal variations, there are several other optical considerations with respect to snow and ice albedo. Snow albedo depends on the angle of incidence of the incoming shortwave radiation. This is a "volume" effect; reflectivity is the net backscatter of radiation, and snow is not a simple surface reflector like a mirror. Rather, it is a granular medium that is partially transparent to shortwave radiation, with complex scattering reactions at granular interstices. The net effect is increased albedo for high zenith angles (low angles of incidence), which occur in early morning and evening. This effect also gives rise to small but measurable differences in the effective surface albedo for cloudy versus clear-sky conditions, as diffuse radiation originates from the full hemisphere whereas the reflection of direct solar radiation is sensitive to zenith angle.

As we saw for thin sea ice, an additional volume effect arises due to the translucency of shallow snowpacks (from millimeters to a few centimeters), as net backscat-ter is influenced by the albedo of underlying layers. The depth of snow as well as contrasts in grain size or crystal geometry in near-surface layers influence the effective albedo of a snowpack.

Sensitivity to the wavelength of the incident radiation adds another important complexity. The broadband albedo is a bulk value for the range of wavelengths corresponding with the solar spectrum. In reality, snow and ice reflectivity vary strongly with wavelength (figure 2.5). Snow and ice are more reflective at visible wavelengths,

Energy Balance Over Snow

Wavelength (pm)

Figure 2.5. Spectral reflectance of snow in the shortwave spectrum, 0.395 to 2.45 |im. From top to bottom, the lines represent grain sizes (effective optical radii) of 50, 200, 500, and 1000 |im. (Data from Anne Nolin, Oregon State University, based on a radiative transfer model.)

Wavelength (pm)

Figure 2.5. Spectral reflectance of snow in the shortwave spectrum, 0.395 to 2.45 |im. From top to bottom, the lines represent grain sizes (effective optical radii) of 50, 200, 500, and 1000 |im. (Data from Anne Nolin, Oregon State University, based on a radiative transfer model.)

with peak reflectivity at 0.46 mm. Snow and ice absorb a higher proportion of radiation in the near-infrared. The effects of grain size, zenith angle, snow depth, and impurities, as described earlier, also vary with wavelength. The effects of impurities and snow depth are strongest at visible wavelengths, whereas grain size exacts its greatest influence at near-infrared wavelengths.

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