Alpine Climate Change and Cryospheric Responses An Introduction


NSIDC/CIRES, University of Colorado, Boulder, CO. 80309-0449, USA


As an introduction to the following chapters dealing with changes in snow and ice conditions in high mountain regions, and their hydrological consequences, a brief overview of recent changes in alpine climates and associated cryospheric responses is presented.

Direct observations and proxy records indicate that historical and recent changes in climate in many mountain regions of the world are at least comparable with, and locally may be greater than, those observed in the adjacent lowlands, Pfister (1985). Actual and potential responses in cryospheric variable include a rise in the snowline, a shorter duration of snow cover, glacier recession, break out of ice-dammed lakes, warming of perennially frozen ground, and thawing of ground ice.

The changes - including the loss of ice core records of climate history as tropical glaciers and ice caps warm and melt water destroys the ice stratigraphy - are of scientific importance. There are also critical socioeconomic implications. These include direct effects of the changes on water resources and hydropower generation, on slope stability, and on hazards relating to avalanches and glacier lakes. Indirect effects include economic and social costs for winter tourism based on skiing and associated sports; and impacts on agricultural, industrial, and consumptive use of water that is strongly influenced by the annual cycle associated with snow and ice melt runoff.


Global mean annual temperature has risen by just over 0.6°C over the last century, with accelerated warming in the last 10 to 15 years. The evidence for changes in climate in mountain areas is both direct and indirect. Observational records are available from the late nineteenth century at a number of mountain observatories, mostly in Europe (Barry 1992). They indicate that mean temperatures have risen by amounts generally comparable with those observed in the lowlands during the twentieth century; however, there are some differences in the pattern of seasonal and diurnal changes. In a survey of available high-elevation data, Diaz and Bradley (1997) present changes in zonally averaged temperatures for 1951-1989 between 30° and 70°N, versus elevation. Mean maximum temperatures increased slightly between 500 and 1500 m, with minor changes at higher elevations, while minimum temperatures rose by about 0.2°C/decade at elevations from 500 m to above 2500 m. In the Rocky mountains, Pepin (2000) documents altitudinal differences in the changes in the Colorado Front Range since 1952, with overall cooling at 3750m but warming between 2500 and 3100m. This results in complex changes in lapse rate. In the tropical Andes, mean annual temperature trends have been determined for 268 stations between 1°N and 23°S, for 1939-1998 (Vuille and Bradley 2000). They find an overall warming of about 0.1°C/decade, but the rate tripled to +0.32-0.34°C/decade over the last

Climate and Hydrology in Mountain Areas. Edited by C. de Jong, D. Collins and R. Ranzi © 2005 John Wiley & Sons, Ltd

25 years. The warming varies with altitude, but there is generally reduced warming with elevation. This is especially apparent on the western (Pacific) slopes of the Andes.

Brown etal. (1992) demonstrated that lapse rates between the high plains (1200 -1500 m) and three stations at 3200 m in the Colorado Rocky mountains had weakened in the daytime, but strengthened at night. Globally, the decrease in diurnal temperature range is attributed to increased cloud cover, locally augmented by changes in precipitation and soil moisture (Dai et al. 1999). An analysis of lapse rates in the Pennines of northern England indicates that atmospheric temperature and moisture level, cloudiness/solar radiation, and wind speed determine lapse rates (Pepin et al. 1999). Thus, changes in lapse rate are complex and may result solely or partly from changes in the frequency of cyclonic/anticyclonic circulation regimes. A shallower/steeper lapse rate may be expected under warmer, moister atmospheric conditions/increased solar radiation. The amplitude of diurnal change in lapse rate intensifies under anticyclonic conditions and slack pressure gradients.

In some mountain regions, monitoring of ground temperatures has begun recently. In the northern Tien Shan, permafrost ground temperatures have risen by 0.2-0.3°C over the last 25 years (Gorbunov et al. 2000). The depth of seasonal freezing has not changed significantly in the low mountains, but there has been a decrease in the depth between 1400 and 2700 m, while above 3000 m the depth of seasonal freezing is increasing. In the Swiss Alps, Haeberli (1994) estimated permafrost warming by about 1°C between 1880 and 1950, then stabilizing, before accelerated warming in the late 1980s to at least 1992. However, a 10-year borehole record (Vonder Muhll etal. 1998) indicates that warming until 1994 was largely compensated by rapid cooling between 1994 and 1996.

Proxy evidence of climatic change is available from changes in glacier size dated by lichenometry and carbon-14, from tree-ring series, and from ice cores, inter alia. Numerous accounts from various mountain regions exemplify these results (Luckman 1997; Luckman and Villalba 2001; Solomina 1999; Kaser 1999). These sources become even more important in mountain regions that lack direct records, or where these are of short duration, as in the Andes and other tropical regions (Barry and Seimon 2000). Diaz and Graham (1996) reported a rise of 100-150m in the altitude of the freezing level in the atmosphere over the inner tropics (10°N-10°S) between 1970 and 1986; this is correlated with a warming in the sea surface over the eastern tropical Pacific. The characteristics of glacier energy balances in the central

Andean region is addressed by Corripio and Purves (Chapter 3).


The effects of global warming on the cryosphere in mountain areas are most visibly manifested in the shrinkage of mountain glaciers and in reduced snow cover duration. However, the responses are by no means linear. For example, warmer winters imply higher atmospheric moisture content and more snowfall is associated with an overall increase in precipitation. Records of glacier length and mass balance during the second half of the twentieth century show reductions in continental climatic regimes, but increases in maritime regimes, such as Norway, southern Alaska and coastal areas of the Pacific Northwest in Canada, and the United States. In the Tropics, the rise in freezing level noted above, as well as changes in atmospheric humidity and perhaps cloudiness, in some cases, has given rise to progressive reduction in mountain glaciers and ice caps over the last century. Particularly, dramatic changes are evident in East Africa where there has been a 75% reduction in ice area on Mount Kilimanjaro since 1912 (Hastenrath and Greischar 1997). The ice cover on East African summits will be lost within 20 years or so, unless there is a dramatic shift in climatic conditions.

In an example of subtle changes in snow cover, Bohm (1986) reported a reduction in May-September snow cover at Sonnblick (3106m), Austria, from 82 days during 1910-1925 to 53 days in 1955-1970. The mean summer temperature was about 0.5°C higher in the second interval. However, the associated change in snow cover duration estimated from average gradients of snow cover duration and temperature lapse rate would only be about 10-11 days (Barry 1990). Such nonlinear responses may arise through local albedo-temperature feedback effects, but this still requires thorough investigation. Keller and Goyette (Chapter 19) provide scenarios of snowmelt in the Swiss Alps under climatic changes.

Large responses are expected in the annual hydrologic regime of rivers where a significant proportion of the runoff is from melt of snow cover and from wastage of ice in heavily glacierized basins. Runoff models under global warming scenarios project a higher and earlier peak of spring runoff from snowmelt and reduced flow in summer (Rango and Martinec 1998). For the upper Rhone, for example, Collins (1987) found discharge correlated with mean summer temperature; a 1°C cooling between 1941-1950 and 1968-1977 led to a 26% decrease in mean summer discharge. Warming trends will have the opposite effect, but a dominant component of runoff change in heavily glacierized basins is attributable to the reduction in ice area. Chen and Ohmura (1990) calculated an 11% decrease in runoff from a basin of the upper Rhone drainage with 66% ice cover between 1922-1929 and 1968-1972, compared with only 6% decrease in one with about 17% ice cover between 1910 -1919 and 1968 -1972. In the latter case, the Rhone at Porte du Scex, runoff changes responded also to changes in climate but a decrease in basin precipitation was offset by the effect of warmer summers increasing the ice melt. The introductory chapter and Chapter 18 address this topic using more recent and extensive data.


Socioeconomic effects of changes in mountain snow and ice characteristics will be both direct and indirect. Direct effects associated with a shorter snow season and shallower snow cover will include the reduction or loss of winter sports facilities, or the necessity for enhanced reliance on snowmaking capabilities, with attendant losses of income and adaptation costs. For the Austrian Alps, losses will be exacerbated at lower elevations. Secondary effects resulting from this change may include the loss of related service activities and income at mountain resorts. Summer tourism may also be affected as scenic mountain glaciers shrink and waste away. Maintaining tourist access to the terminus of the Upper Grindelwald glacier, in retreat since the mid-1980s, for example, has necessitated the construction of a wooden stairway.

The changes in snowmelt runoff and its timing will have direct impacts on hydropower generation and impose requirements for alternative power sources. Power outages and loss of revenue by utility companies may be expected, depending upon the relative contribution ofhydropowerto total electricity generation. In adjacent lowland areas where spring runoff is a major source of water for irrigation and for stocking reservoirs, there may be even greater economic consequences. Changes in snow pack will also affect soil moisture levels in spring and summer, with implications for soil biota, fire risk, and the productivity of mountain pastures and forests (Price and Barry 1997).


Barry, R.G. 1990. Changes in mountain climate and glacio-

hydrological responses. Mt. Res. Dev. 10: 161-70. Barry, R.G. 1992. Mountain climatology and past and potential future climatic changes in mountain regions: A review. Mt. Res. Dev. 12:71-86.

Barry, R.G. and Seimon, A. 2000. Research for mountain area development: Climate fluctuations in the mountains of the Americas and their significance. Ambio 29: 364-70, Corrigendum. Ambio 30, 69.

Bohm, R. 1986. Der Sonnblick. Die 100-Jahrige Geschichte des Observatoriuns und Seiner Forschungstatigkeit, Oesterreichischer Bunderverlag, Vienna, p. 224.

Brown, T.J., Barry, R.G. and Doesken, N.J. 1992. An exploratory study of temperature trends for paired mountain - plains stations in Colorado, Sixth Conference on Mountain Meteorology, American Meteorological Society, Boston, MA, pp. 181 -84.

Chen, J.-Y. and Ohmura, A. 1990. On the influence of Alpine glaciers on runoff. In H. Lang and A. Musy, eds., Hydrology in Mountain Regions I. Hydrological Measurements, The Water Cycle, IAHS Publication No. 193, IAHS Press, Wallingford, CT,pp. 117-25.

Collins, D.N. 1987. Climatic fluctuations and runoff from glacierized alpine basins. In S.L. Solomon, M. Beran and W. Hogg, eds., The influence of climatic change and climatic variability on the hydrological regime and water resources. International Association of Hydrology Publication 168, IAHS Press, Wallingford, UK, pp. 77-89.

Dai, A., Trenberth, K.E. and Karl, T.R. 1999. Effects of clouds, soil moisture, precipitation and water vapor on diurnal temperature range. J. Clim. 12: 2451-73.

Diaz, H.F. and Bradley, R.S. 1997. Temperature variations during the last century at high elevation sites. Clim. Change 36: 253-80.

Diaz, H.F. and Graham, N.E. 1996. Recent changes of tropical freezing heights and the role of sea surface temperature. Nature 383: 152-55.

Gorbunov, A.P., Marchenko, S.S. and Seversky, E.V. 2000. Permafrost and seasonally frozen ground response to climate changes in the northern Tien Shan. Krisfera Zemli 4: 11-17.

Haeberli, W. 1994. Accelerated glacier and permafrost changes in the Alps. In M. Beniston, ed., Mountain Environments in Changing Climates, Routledge, London, pp. 91 -107.

Hastenrath, S. and Greischar, L. 1997. Glacier recession on Kilimanjaro, East Africa, 1912-89. J. Glaciol. 43(145): 4655-59.

Kaser, G. 1999. A review of the fluctuations of modern tropical glaciers. Global Planet. Change 23: 93-103.

Luckman, B.H. 1997. Developing a proxy climate record for the last 300 years in the Canadian Rockies - some problems and opportunities. Clim. Change 36: 455-76.

Luckman, B.H. and Villalba, R. 2001. Assessing the synchrone-ity of glacier fluctuations in the western cordillera of the Americas during the last millennium. In V. Markgraf, ed., Interhemispheric Climate Linkages, Academic Press, San Diego, CA, pp. 119-40.

Pepin, N. 2000. Twentieth century change in the Front Range climate record. Arct. Antarct. Alp. Res. 32: 135-46.

Pepin, N., Benham, D. and Taylor, K. 1999. Modeling lapse rates in the maritime uplands of northern England: Implications for climate change. Arct. Antarct. Alp. Res. 31: 151-64.

Pfister, C. 1985. Snow cover, snowlines and glaciers in central Europe since the sixteenth century. In M.J. Tooley and G.M. Sheail, eds., The Climate Scene, George Allen and Unwin, London, pp. 155-74.

Price, M.F. and Barry, R.G. 1997. Climate change. In B. Messerli and J.D. Ives, eds., Mountains of the World. A Global Priority, Parthenon Publishing, New York, pp. 409-45.

Rango, A.S. and Martinec, J. 1998. Effects of global warming on runoff in mountain basins representing different climatic zones. In H. Weater and C. Kirby, eds., Hydrology in a Changing Environment, Vol. 1, John Wiley, Chichester, pp. 133-39.

Solomina, O. 1999. Gornoe Oledenenie Severnoi Evrazi v Golotsene (Mountain Glaciation in Northern Eurasia During the Holocene), Nauchny Mir, Moscow, p. 263.

VonderMiihll, D.S., Stucki, T. andHaeberli, W. 1998. Borehole temperatures in alpine permafrost: A ten-year series. In A.G. Lewcowitz and M. Allard, eds., Proceedings, The 7th International Permafrost Conference, University of Laval, Quebec, pp. 1089-95.

Vuille, M. and Bradley, R.S. 2000. Mean annual temperature trends and their vertical structure in the tropical Andes. Geophys. Res. Lett. 27: 3885-88.


Use of Positive Degree-Day Methods for Calculating Snow and Ice Melting and Discharge in Glacierized Basins in the Langtang Valley, Central Nepal

RIJAN B. KAYASTHA, YUTAKA AGETA AND KOJI FUJITA Dept. ofHydrospheric-Atmospheric Science, Graduate School of Environmental Studies, Nagoya University, Nagoya 464-8601, Japan


Prediction of melting of snow and ice in a glacierized basin is very important to estimate basin discharge. It is more important in the Himalayas where direct field observations are very difficult to carry out because of rugged and remote mountain terrain. The most important energy source for glacier ablation in the Himalayas is radiation. Many studies have shown that net radiation is the dominant energy source for ablation. The net radiation contributes more than 80% of the total energy supply for ablation in the Nepalese Himalayas (Ohata and Higuchi 1980; Kayastha etal. 1999; Kayastha 2001).

Several models and empirical relations have been proposed to calculate glacier ablation in the Nepalese Himalayas, for example, empirical relations to calculate glacier ablation byAgeta and Higuchi (1984), a simplified model for estimating glacier ablation under debris layer by Nakawo and Takahashi (1982) and Rana etal. (1996), and energy balance modeling for glacier mass balance on Glacier AX010 by Kayastha et al. (1999).

The ablation areas of many glaciers in the Himalayas are covered with debris. Debris has a strong influence on the surface energy balance and melting of the underlying ice. The thermal conductivity (or thermal resistance) and albedo are the main physical characteristics of a debris layer that control heat conduction to the ice-debris interface. Kayastha etal. (2000b) studied the practical prediction of ice melting beneath various thicknesses of debris cover on Khumbu Glacier, Nepal, using a positive degree-day factor. Positive degree-day factors for ablation under various debris thicknesses were found and a practical relationship between debris properties and degree-day factor was established for estimating ablation under a debris layer.

A conceptual runoff model called HYCYMODEL is used in Langtang Khola Basin (Khola means a small river in Nepali) by Fukushima etal. (1991) to estimate streamflow change by global warming. They usedAgeta and Higuchi (1984)'s empirical relation to calculate snow and ice melt. Their study did not take into account the effect of debris on glacier surface, which may accelerate or retard melting of underlying ice depending upon its thickness. Braun etal. (1993) applied the conceptual precipitation-runoff model in the same basin for better understanding of hydrological processes and efficient planning and operation of water resources. Similarly, Rana et al. (1996) used the same

Climate and Hydrology in Mountain Areas. Edited by C. de Jong, D. Collins and R. Ranzi © 2005 John Wiley & Sons, Ltd

HYCYMODEL and empirical relation for melting of snow and ice for modeling runoff from the basin with inclusion of effect of debris on melting of underlying ice. All these three runoff models need daily data of air temperature, precipitation, and other parameters.

Regarding a method to predict snow and ice melt in the Himalayas, the method should be simple with a minimum field data requirement. Therefore, the positive degree-day method is applied to estimate snow and ice melt from debris-free areas as well as ice melt under debris layers. The degree-day method is based on the assumption that the melting of snow or ice during any particular period is proportional to the sum of daily mean temperatures above the melting point during that period, and the sum is called the positive degree-day sum (PDD). The factor linking ablation to this temperature sum is the positive degree-day factor. The degree-day factor involves a simplification of complex processes that are properly described by the energy balance of the glacier surface and overlaying atmospheric boundary layer (Braithwaite and Olesen 1989). This is because the factors that determine the melt process are correlated with temperature or, in other words, the air temperature contains information on the major energy sources. For example, in the net radiation, the incoming longwave radiation is the dominant component of incoming heat source for melt at surface, which transfers information of air temperature to surface (Ohmura2001). It is found that under clear sky about 60% of the atmospheric emission is derived from within the first 100 m and 90% from the first 1 km of the atmosphere. When the sky is overcast with the cloud bottom within the first 1 km, more than 90% originates within this layer between the surface and the bottom of the cloud.

Because of its simplicity and reasonably good results, the degree-day concept has been used by many authors. Braithwaite and Olesen (1989) and Reeh (1991) used the degree-day method to calculate melting over the Greenland ice sheet. Laumann and Reeh (1993) and Johannesson etal. (1995) used the degree-day method for estimating melt rates on different glaciers in Iceland, Norway, and Greenland. Hock (1999) found that the classical degree-day method yields a good simulation of the seasonal pattern of discharge from a small glacier in Sweden. Braithwaite and Zhang (2000) used the degree-day model to study sensitivity of mass balance of five Swiss glaciers to temperature changes.

In this study, the so-called classical degree-day method is used to estimate snow and ice melt, but the PDDs are calculated from monthly mean air temperatures using the concept of Normal distribution (Braithwaite 1985). The main purpose of this paper is to estimate annual discharge from Langtang and Lirung Khola Basins by the degree-day method using monthly mean air temperature and monthly total precipitation. The method is tested with measured discharge from July 1985 to June 1986 in Langtang Khola Basin and May to September 1996 in Lirung Khola Basin. The interannual variation of discharge from 1985 to 1999 is then analyzed. This paper is organized with six sections. The study area is introduced in Section 2.2, data in Section 2.3, and the methodology in Section 2.4. Results and discussion are described in Section 2.5 and conclusions are in the last section.


The investigated basins are located in the Langtang valley, approximately 60 km north of Kathmandu, Nepal. Figure 2.1 shows the location and drainage basins of Langtang Khola and Lirung Khola with hydrological observation sites (S1 and S2) and a meteorological observation site (BH) at an altitude of3920 m a.s.l. The main physical characteristics of the investigated basins are shown in Table 2.1. The altitudinal distribution of Lirung Khola Basin in every 200 m as shown in Rana et al. (1996) is used in this study. In the case of Langtang Khola Basin, snow and ice melt are calculated in every 250 m altitude bands by dividing the drainage basin of 500 m altitude bands as shown in Fukushima et al. (1987) into two equal parts.


Hydrological data used in this study for verifying the calculated discharge are the discharges measured during the hydrological and meteorological observations carried out in Langtang Khola Basin for a full year from July 1985 to June 1986 and from May to September 1996 in Lirung Khola Basin by a joint research team of Japanese and Nepalese scientists. The mean air temperature and total precipitation from July 1985 to June 1986 was 2.7°C and 1225 mm, respectively, at BH. The observed discharge showed that it was mostly concentrated in the period from June to September, coinciding with the summer monsoon period in Nepal. The total observed specific discharge during the above period at S1 was 1358 mm (Fukushima etal. 1987). The monthly mean air temperature and monthly total precipitation from 1988 to 1999 are used to estimate the interannual variation of discharge in this study. These were observed at Kyangjing (3920 m a.s.l.), Langtang hydrometeorological observation station (same area of BH) of Department of Hydrology and Meteorology, His Majesty's Government of Nepal.

Khola Basins. S1, S2 and BH represent hydrological observation sites in Langtang Khola Basin, Lirung Khola Basin and Base House for meteorological observations, respectively
Table 2.1 Main physical characteristics of the investigated basins

Name of the basin

Langtang Khola

Lirung Khola

Name of the area

Langtang valley

Langtang valley

Mountain range



Elevation range of the basin (m a.s.l.)



Elevation range of experimental sites (m a.s.l.)




28°08' -28° 23'N

28°13' -28°16'N


85°35'-85° 48'E


Area (km2)






Glaciers and permanent snow (%)



Dominant vegetation type

No vegetation

No vegetation

Forest (%)



Mean runoff at the catchment outlet (mm)



Mean precipitation (mm)



Monthly mean air temperature (°C)

Figure 2.2 Calculated monthly snowfall amount in precipitation versus monthly mean air temperature on Glacier AX010 from June to August in 1978

Monthly mean air temperature (°C)

Figure 2.2 Calculated monthly snowfall amount in precipitation versus monthly mean air temperature on Glacier AX010 from June to August in 1978

The tendency of precipitation to increase with altitude is seen in glacier areas in the Nepalese Himalayas (Higuchi et al. 1982). In the Langtang valley, the precipitation at 5000 m altitude was 1.3 times larger than at 4000 m in rainy season (Seko 1987). From this observed result, the precipitation was assumed as a function of altitude as follows, since we have precipitation data only at BH.

where anm has either a value of unity or zero according to

If the temperature is assumed to constitute a stationary random series, the time summation in Equation (2.2) can be replaced by an ensemble-summation as follows.


Estimation of snowfall amount during precipitation event is carried out using the relation obtained on Glacier AX010, east Nepal (Figure 2.2). The relation was obtained by plotting calculated monthly snowfall amount in precipitation versus monthly mean air temperature (Kayastha etal. 1999). The mean temperature lapse rate with altitude at Lirung Glacier/BH and Yala Glacier/BH (5.3°Ckm-1 in Fujita etal. 1997) is used to derive the temperature at higher altitudes in both basins.


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