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利用者:Halowand/Hon'yaku

Overview of the Cryosphere and its larger components, from the UN Environment Programme Global Outlook for Ice and Snow.

雪氷圏(cryosphere)とは、地球の表面において水が固体の状況で存在する場所（海氷、湖の氷、川の氷、雪、氷河、氷冠、氷床、永久凍土）をまとめていう言葉。よって氷雪圏は水圏とかなりかぶる。雲、降水、hydrology, そして大気と海水の循環など、表面エネルギーと水分の移動を通じて、氷雪圏は地球全体の気候システムと重大な連携・フィードバックを持ち、必要不可欠な部分になっている。このフィードバックを通じて、氷雪圏は地球全体の気候と気候モデルの全地球的変化に対して大きく影響している。

構造

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氷は地球表面においては主に雪、湖や川の氷、海氷、氷河、氷床、凍土と永久凍土などとして存在する。これらの形で水が存在できる時間にはかなり幅がある。雪と氷は本質的にワンシーズンしか存在できない。北極海の中央を除くほとんどの海氷は、たとえ夏を越せたとしても二三年で溶けてしまう。しかし氷河、氷床、地中氷は10年から100000年以上は凍っていられる。さらに東南極の深いところにある氷は100万年間凍っている可能性がある。

世界にある氷の殆どは南極大陸に、専ら東南極氷床にある。面積だけで言えば北半球の冬に雪と氷が覆う面積が最も広く、1月には平均で北半球の23％を覆う。The large areal extent and the important climatic roles of snow and ice, related to their unique physical properties, indicate that the ability to observe and model snow and ice-cover extent, thickness, and physical properties (radiative and thermal properties) is of particular significance for climate research.

There are several fundamental physical properties of snow and ice that modulate energy exchanges between the surface and the atmosphere. The most important properties are the surface reflectance (albedo), the ability to transfer heat (thermal diffusivity), and the ability to change state (latent heat). These physical properties, together with surface roughness, emissivity, and dielectric characteristics, have important implications for observing snow and ice from space. For example, surface roughness is often the dominant factor determining the strength of radar backscatter .[1] Physical properties such as crystal structure, density, length, and liquid-water content are important factors affecting the transfers of heat and water and the scattering of microwave energy.

The surface reflectance of incoming solar radiation is important for the surface energy balance (SEB). It is the ratio of reflected to incident solar radiation, commonly referred to as albedo. Climatologists are primarily interested in albedo integrated over the shortwave portion of the electromagnetic spectrum (~300 to 3500 nm), which coincides with the main solar energy input. Typically, albedo values for non-melting snow-covered surfaces are high (~80-90%) except in the case of forests. The higher albedos for snow and ice cause rapid shifts in surface reflectivity in autumn and spring in high latitudes, but the overall climatic significance of this increase is spatially and temporally modulated by cloud cover. (Planetary albedo is determined principally by cloud cover, and by the small amount of total solar radiation received in high latitudes during winter months.) Summer and autumn are times of high-average cloudiness over the Arctic Ocean so the albedo feedback associated with the large seasonal changes in sea-ice extent is greatly reduced. Groisman et al. (1994a) observed that snow cover exhibited the greatest influence on the Earth radiative balance in the spring (April to May) period when incoming solar radiation was greatest over snow-covered areas.[2]

The thermal properties of cryospheric elements also have important climatic consequences. Snow and ice have much lower thermal diffusivities than air. Thermal diffusivity is a measure of the speed at which temperature waves can penetrate a substance. Snow and ice are many orders of magnitude less efficient at diffusing heat than air. Snow cover insulates the ground surface, and sea ice insulates the underlying ocean, decoupling the surface-atmosphere interface with respect to both heat and moisture fluxes. The flux of moisture from a water surface is eliminated by even a thin skin of ice, whereas the flux of heat through thin ice continues to be substantial until it attains a thickness in excess of 30 to 40 cm. However, even a small amount of snow on top of the ice will dramatically reduce the heat flux and slow down the rate of ice growth. The insulating effect of snow also has major implications for the hydrological cycle. In non-permafrost regions, the insulating effect of snow is such that only near-surface ground freezes and deep-water drainage is uninterrupted.[3]

While snow and ice act to insulate the surface from large energy losses in winter, they also act to retard warming in the spring and summer because of the large amount of energy required to melt ice (the latent heat of fusion, 3.34 x 105 J/kg at 0°C). However, the strong static stability of the atmosphere over areas of extensive snow or ice tends to confine the immediate cooling effect to a relatively shallow layer, so that associated atmospheric anomalies are usually short-lived and local to regional in scale.[4] In some areas of the world such as Eurasia, however, the cooling associated with a heavy snowpack and moist spring soils is known to play a role in modulating the summer monsoon circulation.[5] Gutzler and Preston (1997) recently presented evidence for a similar snow-summer circulation feedback over the southwestern United States.[6]

The role of snow cover in modulating the monsoon is just one example of a short-term cryosphere-climate feedback involving the land surface and the atmosphere. From Figure 1 it can be seen that there are numerous cryosphere-climate feedbacks in the global climate system. These operate over a wide range of spatial and temporal scales from local seasonal cooling of air temperatures to hemispheric-scale variations in ice sheets over time-scales of thousands of years. The feedback mechanisms involved are often complex and incompletely understood. For example, Curry et al. (1995) showed that the so-called “simple” sea ice-albedo feedback involved complex interactions with lead fraction, melt ponds, ice thickness, snow cover, and sea-ice extent.

雪

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雪は氷雪圏の内では二番目に広い範囲を覆っており、面積は最大でおよそ4700万平方キロメートル。 地球の雪で覆われている地域(SCA)の殆どは北半球に位置している。時間に伴って起きる増減のほとんどは季節の変化によるものであり、北半球では1月に4650万平方キロメートル、8月に380万平方キロメートルになる[7]。北アメリカ大陸の冬に雪で覆われている地域は実近100年では増加傾向にあり、(Brown and Goodison 1996; Hughes et al. 1996)降水量が増えたことがおおよその原因である[8]。しかしながら、the available 衛星画像 show that 北半球の冬 snow cover has exhibited little interannual variability over the 1972-1996 period, with a coefficient of variation (COV=s.d./mean) for January Northern Hemisphere snow cover of < 0.04. According to Groisman et al. (1994a) Northern Hemisphere spring snow cover should exhibit a decreasing trend to explain an observed increase in Northern Hemisphere spring air temperatures this century. Preliminary estimates of SCA from historical and reconstructed in situ snow-cover data suggest this is the case for Eurasia, but not for North America, where spring snow cover has remained close to current levels over most of this century.[9] Because of the close relationship observed between hemispheric air temperature and snow-cover extent over the period of satellite data (IPCC 1996), there is considerable interest in monitoring Northern Hemisphere snow-cover extent for detecting and monitoring climate change.

雪は水収支において猛烈にに重要な保存形態である。特に世界の山岳地帯の季節的な積雪。面積が限られているとはいえ、地球の山岳地帯に降った季節的な積雪は地上を流れる水や地下水の主要源を占めており、中緯度地帯を広く潤している。例えばコロラド川流域から流れ出る一年間の流出分の内85%以上が雪融け水である。 地球の山から流れ出る融雪け水は帯水層と川を潤し、流域に住む人間の水資源になっている。Further, over 40% of the world’s protected areas are in mountains, attesting to their value both as unique ecosystems needing protection and as recreation areas for humans. Climate warming is expected to result in major changes to the partitioning of snow and rainfall, and to the timing of snowmelt, which will have important implications for water use and management. These changes also involve potentially important decadal and longer time-scale feedbacks to the climate system through temporal and spatial changes in soil moisture and runoff to the oceans.(Walsh 1995). Freshwater fluxes from the snow cover into the marine environment may be important, as the total flux is probably of the same magnitude as desalinated ridging and rubble areas of sea ice.[10] In addition, there is an associated pulse of precipitated pollutants which accumulate over the Arctic winter in snowfall and are released into the ocean upon ablation of the sea-ice.

海氷

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海氷 covers much of 極洋 and forms by freezing of海水。1970年代初頭から収集されていた人工衛星データにより、南北半球を覆う海氷は季節的、地域的、各年的にかなり変動することがわかった。季節的には、南半球の海氷の面積は5倍ちかく変動し、最小量は2月の300-400万平方キロメートル、最大量は9月の1700-2000万平方キロメートルである[11][12]。

北半球の極洋、すなわち北極海では季節的な変動は小さい。ここは他地域とつながりが少なく、かつ高緯度であるため、広い範囲を一年中氷が覆っている。また陸地が取り囲んでいるため、冬に海氷が赤道側へ氷が広がるにも限界がある。よって季節的には、北半球での海氷面積は2倍しか変動せず、最小量は9月の700-900万平方キロメートルで最大量は3月の1400-1600万平方キロメートルである[12][13]。

The ice cover exhibits much greater regional-scale interannual variability than it does hemispherical. For instance, in the region of the Seas of Okhotsk and Japan, maximum ice extent decreased from 1.3 million km² in 1983 to 0.85 million km² in 1984, a decrease of 35%, before rebounding the following year to 1.2 million km² .[12] The regional fluctuations in both hemispheres are such that for any several-year period of the satellite record some regions exhibit decreasing ice coverage while others exhibit increasing ice cover.[14] The overall trend indicated in the passive microwave record from 1978 through mid-1995 shows that the extent of Arctic sea ice is decreasing 2.7% per decade.[15] Subsequent work with the satellite passive-microwave data indicates that from late October 1978 through the end of 1996 the extent of Arctic sea ice decreased by 2.9% per decade while the extent of Antarctic sea ice increased by 1.3% per decade.[16]

Lake ice and river ice

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Ice forms on rivers and lakes in response to seasonal cooling. The sizes of the ice bodies involved are too small to exert other than localized climatic effects. However, the freeze-up/break-up processes respond to large-scale and local weather factors, such that considerable interannual variability exists in the dates of appearance and disappearance of the ice. Long series of lake-ice observations can serve as a proxy climate record, and the monitoring of freeze-up and break-up trends may provide a convenient integrated and seasonally specific index of climatic perturbations. Information on river-ice conditions is less useful as a climatic proxy because ice formation is strongly dependent on river-flow regime, which is affected by precipitation, snow melt, and watershed runoff as well as being subject to human interference that directly modifies channel flow, or that indirectly affects the runoff via land-use practices.

Lake freeze-up depends on the heat storage in the lake and therefore on its depth, the rate and temperature of any inflow, and water-air energy fluxes. Information on lake depth is often unavailable, although some indication of the depth of shallow lakes in the Arctic can be obtained from airborne radar imagery during late winter (Sellman et al. 1975) and spaceborne optical imagery during summer (Duguay and Lafleur 1997). The timing of breakup is modified by snow depth on the ice as well as by ice thickness and freshwater inflow.

Frozen ground and permafrost

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Frozen ground (permafrost and seasonally frozen ground) occupies approximately 54 million km² of the exposed land areas of the Northern Hemisphere (Zhang et al., 2003) and therefore has the largest areal extent of any component of the cryosphere. Permafrost (perennially frozen ground) may occur where mean annual air temperatures (MAAT) are less than -1 or -2°C and is generally continuous where MAAT are less than -7°C. In addition, its extent and thickness are affected by ground moisture content, vegetation cover, winter snow depth, and aspect. The global extent of permafrost is still not completely known, but it underlies approximately 20% of Northern Hemisphere land areas. Thicknesses exceed 600 m along the Arctic coast of northeastern Siberia and Alaska, but, toward the margins, permafrost becomes thinner and horizontally discontinuous. The marginal zones will be more immediately subject to any melting caused by a warming trend. Most of the presently existing permafrost formed during previous colder conditions and is therefore relic. However, permafrost may form under present-day polar climates where glaciers retreat or land emergence exposes unfrozen ground. Washburn (1973) concluded that most continuous permafrost is in balance with the present climate at its upper surface, but changes at the base depend on the present climate and geothermal heat flow; in contrast, most discontinuous permafrost is probably unstable or "in such delicate equilibrium that the slightest climatic or surface change will have drastic disequilibrium effects".[17]

Under warming conditions, the increasing depth of the summer active layer has significant impacts on the hydrologic and geomorphic regimes. Thawing and retreat of permafrost have been reported in the upper Mackenzie Valley and along the southern margin of its occurrence in Manitoba, but such observations are not readily quantified and generalized. Based on average latitudinal gradients of air temperature, an average northward displacement of the southern permafrost boundary by 50-to-150 km could be expected, under equilibrium conditions, for a 1°C warming.

Only a fraction of the permafrost zone consists of actual ground ice. The remainder (dry permafrost) is simply soil or rock at subfreezing temperatures. The ice volume is generally greatest in the uppermost permafrost layers and mainly comprises pore and segregated ice in Earth material. Measurements of bore-hole temperatures in permafrost can be used as indicators of net changes in temperature regime. Gold and Lachenbruch (1973) infer a 2-4°C warming over 75 to 100 years at Cape Thompson, Alaska, where the upper 25% of the 400-m thick permafrost is unstable with respect to an equilibrium profile of temperature with depth (for the present mean annual surface temperature of -5°C). Maritime influences may have biased this estimate, however. At Prudhoe Bay similar data imply a 1.8°C warming over the last 100 years (Lachenbruch et al. 1982). Further complications may be introduced by changes in snow-cover depths and the natural or artificial disturbance of the surface vegetation.

The potential rates of permafrost thawing have been established by Osterkamp (1984) to be two centuries or less for 25-meter-thick permafrost in the discontinuous zone of interior Alaska, assuming warming from -0.4 to 0°C in 3–4 years, followed by a further 2.6°C rise. Although the response of permafrost (depth) to temperature change is typically a very slow process (Osterkamp 1984; Koster 1993), there is ample evidence for the fact that the active layer thickness quickly responds to a temperature change (Kane et al. 1991). Whether, under a warming or cooling scenario, global climate change will have a significant effect on the duration of frost-free periods in both regions with seasonally- and perennially-frozen ground.

氷河と氷床

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氷河と氷床は固体の台地の上を流れる氷の塊である。雪の堆積、表面と基底部の融解、取り囲む海や湖への崩落、そして氷内部のエネルギーによりコントロールされている。The latter results from gravity-driven creep flow ("glacial flow") within the ice body and sliding on the underlying land, which leads to thinning and horizontal spreading.[18] Any imbalance of this dynamic equilibrium between mass gain, loss and transport due to flow results in either growing or shrinking ice bodies.

氷床は世界最大の淡水源であり、およそ77％がキープされている。氷床のおよそ9割は南極にあり、残りの1割の殆どはグリーンランドにある。残りの氷と氷河は0.5％もない。もしすべて解ければ世界の海面は80メートル上がると言われている。Because of their size in relation to annual rates of snow accumulation and melt, the residence time of water in ice sheets can extend to 10万年から100万年. Consequently, any climatic perturbations produce slow responses, occurring over glacial and interglacial periods. Valley glaciers respond rapidly to climatic fluctuations with typical response times of 10–50 years.[19] However, the response of individual glaciers may be asynchronous to the same climatic forcing because of differences in glacier length, elevation, slope, and speed of motion. Oerlemans (1994) provided evidence of coherent global glacier retreat which could be explained by a linear warming trend of 0.66°C per 100 years.[19]

While glacier variations are likely to have minimal effects upon global climate, their recession may have contributed one third to one half of the observed 20th Century rise in sea level (Meier 1984; IPCC 1996). Furthermore, it is extremely likely that such extensive glacier recession as is currently observed in the Western Cordillera of North America,[20] where runoff from glacierized basins is used for irrigation and hydropower, involves significant hydrological and ecosystem impacts. Effective water-resource planning and impact mitigation in such areas depends upon developing a sophisticated knowledge of the status of glacier ice and the mechanisms that cause it to change. Furthermore, a clear understanding of the mechanisms at work is crucial to interpreting the global-change signals that are contained in the time series of glacier mass balance records.

Combined glacier mass balance estimates of the large ice sheets carry an uncertainty of about 20%. Studies based on estimated snowfall and mass output tend to indicate that the ice sheets are near balance or taking some water out of the oceans.[21] Marinebased studies [22] suggest sea-level rise from the Antarctic or rapid ice-shelf basal melting. Some authors (Paterson 1993; Alley 1997) have suggested that the difference between the observed rate of sea-level rise (roughly 2 mm/y) and the explained rate of sea-level rise from melting of mountain glaciers, thermal expansion of the ocean, etc. (roughly 1 mm/y or less) is similar to the modeled imbalance in the Antarctic (roughly 1 mm/y of sea-level rise; Huybrechts 1990), suggesting a contribution of sea-level rise from the Antarctic.

Relationships between global climate and changes in ice extent are complex. The mass balance of land-based glaciers and ice sheets is determined by the accumulation of snow, mostly in winter, and warm-season ablation due primarily to net radiation and turbulent heat fluxes to melting ice and snow from warm-air advection,[23][24](Munro 1990). However, most of Antarctica never experiences surface melting.[25] Where ice masses terminate in the ocean, iceberg calving is the major contributor to mass loss. In this situation, the ice margin may extend out into deep water as a floating ice shelf, such as that in the Ross Sea. Despite the possibility that global warming could result in losses to the Greenland ice sheet being offset by gains to the Antarctic ice sheet,[26] there is major concern about the possibility of a West Antarctic Ice Sheet collapse. The West Antarctic Ice Sheet is grounded on bedrock below sea level, and its collapse has the potential of raising the world sea level 6–7 m over a few hundred years.

Most of the discharge of the West Antarctic Ice Sheet is via the five major ice streams (faster flowing ice) entering the Ross Ice Shelf, the Rutford Ice Stream entering Ronne-Filchner shelf of the Weddell Sea, and the Thwaites Glacier and Pine Island Glacier entering the Amundsen Ice Shelf. Opinions differ as to the present mass balance of these systems (Bentley 1983, 1985), principally because of the limited data. The West Antarctic Ice Sheet is stable so long as the Ross Ice Shelf is constrained by drag along its lateral boundaries and pinned by local grounding.

関連項目

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出典

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  1. ^ Hall, D. K., 1996: Remote sensing applications to hydrology: imaging radar. Hydrological Sciences, 41, 609-624.
  2. ^ Groisman, P. Ya, T. R. Karl, and R. W. Knight, 1994a: Observed impact of snow cover on the heat balance and the rise of continental spring temperatures. Science, 363, 198-200.
  3. ^ Lynch-Stieglitz, M., 1994: The development and validation of a simple snow model for the GISS GCM. J. Climate, 7, 1842-1855.
  4. ^ Cohen, J., and D. Rind, 1991: The effect of snow cover on the climate. J. Climate, 4, 689-706.
  5. ^ Vernekar, A. D., J. Zhou, and J. Shukla, 1995: The effect of Eurasian snow cover on the Indian monsoon. J. Climate, 8, 248-266.
  6. ^ Gutzler, D. S., and J. W. Preston, 1997: Evidence for a relationship between spring snow cover in North America and summer rainfall in New Mexico. Geophys. Res. Lett., 24, 2207-2210.
  7. ^ Robinson, D. A., K. F. Dewey, and R. R. Heim, 1993: Global snow cover monitoring: an update. Bull. Amer. Meteorol. Soc., 74, 1689-1696.
  8. ^ Groisman, P. Ya, and D. R. Easterling, 1994: Variability and trends of total precipitation and snowfall over the United States and Canada. J. Climate, 7, 184-205.
  9. ^ Brown, R. D., 1997: Historical variability in Northern Hemisphere spring snow covered area. Annals of Glaciology, 25 (in press). Axel Heiberg Island, N.W.T., Canada, 1960-91. J. Glaciology, 42(142): 548-563.
  10. ^ Prinsenberg, S. J. 1988: Ice-cover and ice-ridge contributions to the freshwater contents of Hudson Bay and Foxe Basin. Arctic, 41, 6-11.
  11. ^ Zwally, H. J., J. C. Comiso, C. L. Parkinson, W. J. Campbell, F. D. Carsey, and P. Gloersen, 1983: Antarctic Sea Ice, 1973-1976: Satellite Passive-Microwave Observations. NASA SP-459, National Aeronautics and Space Administration, Washington, D.C., 206 pp.
  12. ^ a b c Gloersen, P., W. J. Campbell, D. J. Cavalieri, J. C. Comiso, C. L. Parkinson, and H. J. Zwally, 1992: Arctic and Antarctic Sea Ice, 1978-1987: Satellite Passive-Microwave Observations and Analysis. NASA SP-511, National Aeronautics and Space Administration, Washington, D.C., 290 pp.
  13. ^ Parkinson, C. L., J. C. Comiso, H. J. Zwally, D. J. Cavalieri, P. Gloersen, and W. J. Campbell, 1987: Arctic Sea Ice, 1973-1976: Satellite Passive-Microwave Observations, NASA SP-489, National Aeronautics and Space Administration, Washington, D.C., 296 pp.
  14. ^ Parkinson, C. L., 1995: Recent sea-ice advances in Baffin Bay/Davis Strait and retreats in the Bellinshausen Sea. Annals of Glaciology, 21, 348-352.
  15. ^ Johannessen, O. M., M. Miles, and E. Bjørgo, 1995: The Arctic’s shrinking sea ice. Nature, 376, 126-127.
  16. ^ Cavalieri, D. J., P. Gloersen, C. L. Parkinson, J. C. Comiso, and H. J. Zwally, 1997: Observed hemispheric asymmetry in global sea ice changes. Science, 278, 1104-1106.
  17. ^ Washburn, A. L., 1973: Periglacial processes and environments. Edward Arnold, London, 320 pp. p.48
  18. ^ Greve, R.; Blatter, H. (2009). Dynamics of Ice Sheets and Glaciers. Springer. doi:10.1007/978-3-642-03415-2. ISBN 978-3-642-03414-5 
  19. ^ a b Oerlemans, J., 1994: Quantifying global warming from the retreat of glaciers. Science, 264, 243-245.
  20. ^ Pelto, M. S., 1996: Annual net balance of North Cascade Glaciers, 1984-94. J. Glaciology, 42, 3-9.
  21. ^ Bentley, C. R., and M. B. Giovinetto, 1991: Mass balance of Antarctica and sea level change. In: G. Weller, C. L. Wilson and B. A. B. Severin (eds.), Polar regions and climate change. University of Alaska, Fairbanks, p. 481-488.
  22. ^ Jacobs, S. S., H. H. Helmer, C. S. M. Doake, A. Jenkins, and R. M. Frohlich, 1992: Melting of ice shelves and the mass balance of Antarctica. J. Glaciology, 38, 375-387.
  23. ^ Paterson, W. S. B., 1993: World sea level and the present mass balance of the Antarctic ice sheet. In: W.R. Peltier (ed.), Ice in the Climate System, NATO ASI Series, I12, Springer-Verlag, Berlin, 131-140.
  24. ^ Van den Broeke, M. R., 1996: The atmospheric boundary layer over ice sheets and glaciers. Utrecht, Universitiet Utrecht, 178 pp..
  25. ^ Van den Broeke, M. R., and R. Bintanja, 1995: The interaction of katabatic wind and the formation of blue ice areas in East Antarctica. J. Glaciology, 41, 395-407
  26. ^ Ohmura, A., M. Wild, and L. Bengtsson, 1996: A possible change in mass balance of the Greenland and Antarctic ice sheets in the coming century. J. Climate, 9, 2124-2135.

外部リンク

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