1 Snowcover Structure and Metamorphism Snow stratification results from successive snowfalls over the winter and processes that transform the snow cover between snowfalls. Snow metamorphism depends on temperature, temperature gradient, and liquid water content. The size, type, and bonding of snow crystals are responsible for pore size and permeability of the snowpack.

2 In low wind speed environments, fresh snowfall has low hardness and density (50 to 120 kg m -3 ). Temperature gradients induce water vapour pressure gradients, vapour diffusion from the warmest crystals, and consequent change in the shape of the crystals. If persistent temperature profiles exist, then distinctive crystal shapes develop within the snowpack.

3 Source: DeWalle & Rango (2008)

4 First, crystals may be transformed into faceted crystals and eventually into depth hoar over time. The grain shape of these crystals does not allow an efficient compaction of the snowpack. Depth hoar may take a number of shapes including cups, needles, scrolls and plates. A layer of depth hoar has very low strength.

5 Source: DeWalle & Rango (2008)

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7 Metamorphism can also result from compaction caused by the pressure of overlying layers of snow that leads to its densification. This process is responsible for transforming snow into glacial ice whose crystals sometimes attain sizes of the order of 1 cm. During its early stages, the refreezing of melt water can accelerate the densification process. Snow density is often assumed to increase exponentially with time (e.g. Verseghy, 1991).

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9 Source: DeWalle & Rango (2008)

10 Summary of snow metamorphosis processes Source: DeWalle & Rango (2008)

11 Snowpack Properties and Evolution The preceding slides emphasized the various processes of metamorphism that control the snow bulk properties. Thermal properties that depend only on density (specific heat, latent heat) are well defined.

12 However, those that depend on conductivity or permeability of the snowpack are affected by sintering, particle size, ice layers and depth hoar. The specific and latent heats of snow are the simplest thermal properties to determine since the contributions from air and water vapour can be discounted; each property is simply the product of the snow density and the corresponding property for ice.

13 The temperature dependence of the specific heat of ice given by Dorsey (1940) is: C = T where C is the specific heat (kJ kg -1 K -1 ), and T ( o C) is temperature. The latent heat of melting of ice at 0 o C and standard atmospheric pressure is kJ kg -1.

14 For one-dimensional, steady-state heat flow by conduction in a solid the thermal conductivity is the proportionality constant of the Fourier equation: F = -K dT/dz where F is the heat flux (W m -2 ) and dT/dz is the temperature gradient. The thermal conductivity of snow (K) is a more complex property than specific heat because its magnitude depends on such factors as the density, temperature and the microstructure of the snow.

15 The thermal conductivity of ice varies inversely with temperature by about 0.17% o C -1 ; the same may be expected for snow. A temperature gradient could induce a transfer of vapour and the subsequent release of the latent heat of vapourization, thereby changing the thermal conductivity value.

16 In non-aspirated dry snow the heat transfer process involves: conduction of heat in the network of ice grains and bonds, conduction across air spaces or pores, convection and radiation across pores (probably negligible) and vapour diffusion through the pores. Because of the complexity of the heat transfer processes, the thermal conductivity of snow is generally taken to be an “apparent” or “effective” conductivity K e to embrace all the heat transfer processes.

17 The degree of surface packing (for example, hardness) also affects the flow of heat through snow, probably because a surface crust of low air permeability inhibits ventilation in the upper snow layer. The thermal conductivity of snow, even when dense, is very low compared to that of ice or liquid water; therefore snow is a good insulator. This is an important factor affecting heat loss from buildings and the rate of freezing of lake and river ice.

18 Typical numerical models of snow use three prognostic variables to define a snowpack: snow depth, snow water equivalent (SWE), and temperature. From snow depth and snow water equivalent, one can infer the snow density from: ρ s = ρ w (w/s) where w (m) is SWE, s (m) is the snow depth, and ρ s and ρ w are the snow and water densities, respectively.

19 Source: Sun et al. (2004)

20 Apart from snow depth and SWE, the heat content or temperature of the snowpack is required to describe the system completely. The snow temperature is directly related to its heat content H (J) by: T = H/(ρ w w C). The energy balance of a snowpack is complicated not only by the fact that shortwave radiation penetrates the snow but also by water movement and phase changes.

21 Source: Lynch-Stieglitz (1994) Sleepers River watershed, Vermont

22 The energy balance of a snow volume depends upon whether it is a “cold” (< 0 o C) or a “wet” (0 o C, often isothermal) snowpack. Recall the energy balance of the snowpack: Q* + Q P = Q H + Q E + Q G + ΔQ S + Q M. A term is added here to the energy balance to consider the heat transported by precipitation (Q P ), either snowfall or rainfall.

23 In the case of a cold snowpack, such as is commonly found in mid-latitudes during winter with little or no solar input, Q E and Q M are likely to be negligible. Similarly, heat conduction within the snow will be small because of the low thermal conductivity of snow and the lack of solar heating, so that ΔQ S and Q G are also negligible. The energy balance therefore reduces to that between a net radiative sink Q* and a convective sensible Q H heat source.

24 Source: Armstrong & Brun (2008)

25 Although snowcover reduces the available energy at the surface because of its high albedo to solar radiation and high emissivity of longwave radiation, its insulative properties exert the greatest influence on soil temperature regime. Snow acts as an insulating layer that reduces the upward flux of heat, resulting in higher ground temperatures than would occur if the ground was bare.

26 Source: Armstrong & Brun (2008)

27 In Canada, average near-surface soil temperatures are about 3 o C warmer than average air temperatures. In the case of a “wet” snowpack during the melt period, the surface temperature will remain close to 0 o C, but the air temperature may be above freezing. Since snow is porous, liquid water infiltration is also important in transporting energy within the snowpack and into soils.

28 If meltwater freezes within the snowpack, there is latent release, warming snowpack layers to the freezing point. Most of the energy exchanges between snow and its environment occur at the atmosphere or ground interfaces; however, because snow is porous, some radiation and convective fluxes that occur within the top few centimetres of the snowpack.

29 The important fluxes that can directly penetrate the snowpack are radiation, conduction, convection, and meltwater or rainwater percolation. Temperature regimes in dry snowpacks are exceedingly complex and are controlled by a balance of the energy regimes at the top and bottom of the snowpack, radiation penetration, effective thermal conductivity of the snow layers, water vapour transfer, and latent heat exchange during metamorphism.

30 Temperature stratification within dry snowpacks is usually unstable (warm temperatures below cold temperatures) from formation until late winter and spring, as energy inputs from the soil boundary exceed those from the atmosphere and upper layers. As a result, temperatures become warmer with depth, with gradients as high as 50 o C m -1 in shallow subarctic and arctic snowpacks during early midwinter.

31 In cold climates with frozen soils, an inversion can develop in late winter where the upper snowpack warms to higher temperatures than the lower layers (a stable regime). This reflects higher energy inputs from the atmosphere (often due to long sunlit periods in the northern spring) than from the frozen soil. For a given climate, the thermal regime in the snowpack strongly depends on the amount of snowfall early in the winter season.

32 Heavy snowfall early in the winter will tend to maintain the snowpack relatively warm, whereas shallow snowcovers will adjust more rapidly to the air temperatures. For a deep snowpack a midwinter rainfall would increase density and decrease depth. Subarctic and arctic snowpacks can undergo melt in upper layers whilst maintaining snow temperatures significantly below the freezing point in the lower layers.

33 Internal heat fluxes in wet snow, or in partially wet snow, are principally driven by conduction and by latent heat release due to refreezing of liquid water.

34 Ref: Bartelt and Lehning (2002)

35 Ref: Bartelt and Lehning (2002)

36 Source: Pomeroy and Brun (2001)