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Physical Geography of Northern Eurasia

Permafrost

<<< Cryogenic Processes and Relief | Physical Geography Index | Engineering Aspects and Protection of Permafrost >>>

Climate and Permafrost

Climatic Change and Evolution of Permafrost in the Past

Global cooling of climate and perennial freezing of the ground began in the Northern Hemisphere in the late Neogene-early Pleistocene (see above; Baulin and Danilova, 1988; French, 1996). Evolution of the cryolithozone in Russia has been determined not only by climatic changes, but also by cover glaciations and marine transgressions and regressions during the Quaternary (Velichko, 1973; Kudryavtsev et al., 1978b; Baulin and Danilova, 1988). In the western part of the Russian cryolithozone (westwards of the Yenisey), permafrost developed in response to climatic and glacial oscillations in the Quaternary, and marine transgression in the middle Pleistocene and at the beginning of the late Pleistocene. Eastern Siberia has never been glaciated, except for the mountains, which accounts for the most broad extent of permafrost. In the Pleistocene, cooling of climate was accompanied by marine regressions, and arid permafrost conditions in Eastern Siberia were at their extreme (Geocryology of the USSR: Eastern Siberia and the Far East, 1989). The presence of the relict ice wedges confirms a steady permafrost trend. In northeastern Siberia, permafrost has existed uninterruptedly throughout the last 1.5-2 million years, while in Europe and Western Siberia in the Quaternary permafrost developed and thawed repeatedly. There is little evidence on the early stage of permafrost evolution. Only four main stages of Quaternary permafrost variations can be reconstructed: (1) from the early Pleistocene to the beginning of the late Pleistocene; (2) the late Pleistocene; (3) the first half of the Holocene; (4) the late Holocene (Baulin and Danilova, 1988).

The duration of the first stage exceeded 1.5 million years, during which the spread and thickness of permafrost were continuously increasing (Baulin and Danilova, 1988). The first perennial frozen ground was formed in the north of Central and Eastern Siberia and in the Taymyr peninsula. During the early Pleistocene (approximately 2-0.7 millions years ago) permafrost developed across Western Siberia (reaching 54-55 N) and central and western Sakha-Yakutia (Geocryology of the USSR: Western Siberia, 1989). The ice-wedge pseudomorphs and ice-wedge casts of Western Siberia and the relict ice wedges of north-eastern Siberia are testimonies of ancient permafrost. During the glacial period of the middle Pleistocene, transgression of the Arctic Ocean extended to the European north-east and northern Western Siberia, while a continental ice sheet occupied vast areas of Europe (Velichko, 1973; Baulin and Danilova, 1988). Permafrost extended southwards of the glacial sheet and coastline, and the southern limit of permafrost reached 50N in Europe and Kazakhstan.

At the same time, syngenetic deposits with high ice content and thick ice wedges developed in northeastern Siberia and central Sakha-Yakutia (Regional Cryolithology, 1989). During the interglacials of the middle Pleistocene and of the beginning of the late Pleistocene, ice sheets disappeared and permafrost became more limited in extent in Europe and the coastal sector of the Siberian Seas (Velichko, 1973).

At the second stage of its evolution (from 150-90 to 11-10 thousand years ago), permafrost attained the maximum territorial extent, thickness, and the minimum temperatures in Northern Eurasia. On the European and Siberian plains, the southern limit of permafrost extended to 48-49N, where its thickness could attain 300-600 m (Velichko, 1975; Kondratjeva et al., 1993). For the last cold period in Northern Eurasia a magnitude of temperature depression of 5-9C is suggested (Vasilchuk, 1992) while the lowest ground temperatures may have reached -22C to -25C in northern Sakha-Yakutia (Kaplina, 1981). Syngenetic moraine sediments containing ice wedges formed in the nearshore areas of Western Siberia and syngenetic alluvial deposits with ice wedges developed in north-eastern Siberia (Kondratjeva et al., 1993). Europe and Siberia experienced cold non-glacial conditions and were the major areas for the development of the late Pleistocene periglacial and permafrost environments.

The third stage or the Holocene climatic optimum (from 10 to 4.5-3 thousand years ago) was the warmest and driest. Global climatic warming and thawing of permafrost in the contemporary mid-latitudes characterized the Holocene optimum in Eurasia (Baulin and Danilova, 1988). Air temperature increased up to 4.4-5.2C in the high latitudes (Velichko, 1975; Velichko and Nechaev, 1992). The southern limit of permafrost coincided with the Arctic Circle in Western Siberia, ran along 60-61N in Eastern Siberia, diving south and encirled the Transbaikalian mountains. One of the most distinct manifestations of spatial permafrost decrease is the presence of a thaw layer above the second relict layer of the Pleistocene permafrost in the European north-east and Western Siberia. In Western Siberia, the depth of perennial thaw layer increased from 30-50 m at the Arctic Circle to 50-100 m at 64-66N and more than 100-300 m further south (Velichko, 1975; Velichko and Nechaev, 1992). There were two stages of permafrost degradation in Western Siberia (Astakhov, 1995). The first stage began when surface water bodies warmed enough to perforate permafrost by numerous sinking thermokarst ponds and to produce local taliks. The second stage, which only took place to the south of the Arctic Circle, began when climate became so warm that direct solar heating of the deposits was capable of producing regional taliks or a perennially thawed layer separating the Pleistocene permafrost from the seasonally frozen layer. The main evidence of this stage is widespread and varying thermokarst formed in the contemporary cryolithozone and ice-wedge casts in the area of ancient cryolithozone. Such thermokarst forms as the alas plains of northern Sakha-Yakutia developed when the mean summer air temperature was 4-5C higher than at present. Ice-wedge casts formed when the retreat of permafrost was accompanied by melting of ice wedges and infilling of fissures with soil at such a slow rate that the shape of wedges was not destroyed.

The fourth stage, which continued for 3500-4000 years in the late Holocene, is characterized by aggregation of permafrost. The southern boundary of permafrost moved several degrees further south, ground temperatures decreased by 1-2C, and thermal contraction cracking and frost heave intensified (Baulin and Danilova, 1988). The present-day permafrost environment provides numerous examples of late Holocene periglacial modification.

Quaternary permafrost is discussed above.

Global Warming and Permafrost

There are two major approaches to studying the effects of global climatic warming on permafrost: analysis of the long-term meteorological and ground temperature records (Pavlov, 1994, 1997) and modelling (Velichko and Nechaev, 1992; Pavlov, 1997; Ershov, 1997). Recent climatic trends are discussed above. More specifically, the long-term records at Arkhangelsk, Salekhard, and Yakutsk indicate that between 1885 and 1993 mean annual air temperatures have increased by 1.7-2.5C (Pavlov, 1994, 1997). Between 1965 and 1995 an increase of 0.6-1.6C was observed which corresponds to an annual temperature increase of 0.02-0.03C a-1 in northern Europe, 0.03-0.()7C a-1 in Western Siberia, and 0.01-0.08C a-1 in Sakha-Yakutia (Pavlov, 1997). The precipitation and snow cover data indicate that between 1950 and 1980 snow cover depth was increasing at a rate of 0.5-0.7 cm a-1 while winter precipitation was growing, on average, by 1.6 mm a-1 (Pavlov, 1997). More data on the contemporary warming is given above. Combined together, higher temperatures and greater snow depth may result in warming of permafrost. Data, obtained at the permafrost monitoring stations in Western Siberia, show that during the last twenty years the temperatures of the permafrost at a depth of 3 m have increased by 2.0-2.5C and at a depth of 10 m by 1C (Pavlov, 1994, Pavlov, 1997). Results of Global Circulation Models (GCM) predict that high latitudes will warm more than the global mean, particularly in winter (Houghton et al., 1995). The anticipated changes in temperature are considerably higher than the changes indicated by the instrumental records and it may be hypothesized that permafrost environments will be critically affected. This in turn will have a profound influence on climate through feedback mechanisms, such as the additional input of methane and nitrous oxide to the atmosphere (Khalil and Rasmussen, 1989; Kvenvolden, 1993; Kvenvolden and Lorenson, 1993). These may lead to an additional 0.4C increase in global temperatures by 2020 and 0.6-0.7C increase by 2050 (Street and Melnikov, 1990). Various scenarios of climatic change are used to predict permafrost evolution (Borisenkov, 1990; Ershov, 1997). A scenario based upon 4-8C warming predicts that the southern boundary of the cryolithozone will move 200-350 km north in Western Siberia and 350-500 km in the south-west of Eastern Siberia (Ershov, 1997). Between 60 per cent and 70 per cent of the cryolithozone will be occupied by relict permafrost with a thaw layer of 5-20 m. Island and continuous permafrost will occur only in the Asian part of Russia; the continuous permafrost zone will extend across the tundra plains, the northern sector of Central Siberian plateau and the mountains of the north-east. Permafrost temperature will exceed the contemporary temperature by 4-5C and range between +1C and -8C. The rates of cryogenic processes (increase of seasonally thaw layer depth and seasonally frozen layer detachments, thaw slumping, thermokarst and thermoerosion, slope processes) will be similar to those observed during the Holocene climatic optimum (Koster, 1993; French, 1996). However, it should be noted that most recent modelling studies, which use coupled ocean-atmosphere GCMs and incorporate aerosol feedback, suggest more modest rates of climatic warming (Houghton et al., 1995).

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