RUSSIAN GEOGRAPHICAL SOCIETY

FACULTY OF GEOGRAPHY, M.V. LOMONOSOV MOSCOW STATE UNIVERSITY

INSTITUTE OF GEOGRAPHY, RUSSIAN ACADEMY OF SCIENCES

No. 01 [02]

2009

GEOGRAPHY ENVIRONMENT SUSTAINABILITY

Vyacheslav N. Konishchev M.V. Lomonosov Moscow State University, Leninskie gory, 119991, Moscow, Russia E-mail: vkonish@mail.ru

CLIMATE WARMING AND PERMAFROST

ABSTRACT

The purpose of the paper is to discuss re-sponses of the cryolithozone to two different scale warming events: (1) climate warming of recent decades at the end of the 20 t h and the beginning of the 21 s t century, and (2) climate warming in the interval between the cryochron in the late Pleistocene through the Holocene. Changes in several permafrost parameters that occur with climate warm-ing were influenced by the entire complex of landscape characteristics and landscape components that also change under climate warming. This phenomenon may explain why in certain landscapes during climate warming aggradational changes, such as a decrease or stabilization of permafrost temperatures, decrease in the thickness of the seasonal freezing/thawing layer, and increase in the ice content of the upper permafrost layers, are present along with degradation processes (e. g., permafrost temperature increase, increase in the area of taliks, etc.)

KEYWORDS: climatic and non-climatic factor of heat exchange, lateral and frontal permafrost degradation, protective layer, ice complex.

PROBLEM STATEMENT From very early phases of its development, permafrost science has focused on the cryolithozone (permafrost) response to climate change. This problem became especially important in the early 1960s, when the issue of global warming emerged.

To date, dozens of papers and monographs and many international and Russian conferences specifically targeted changes in the cryolithozone resulting from global climate change. For more than 15 years, countries with permafrost or even short-term seasonal episodes of freezing of soils and the upper layers of the lithosphere within their territory have been engaged in several international programs, e. g., Circumpolar Active Layer Monitoring (CALM) and the Thermal Status of Permafrost (TSP).

The collective research effort has established that climate is not the only factor that influences the important characteristics of permafrost, namely, the average annual temperature of frozen soils at the depth of zero annual amplitude (tav), thickness of the seasonal thawing layer (STL), and other parameters. The impact of air temperatures, snow cover, different types of vegetation, soils, lithology, and water content on these permafrost parameters, other properties of frozen soils, and cryogenic processes (thermokarst, thermal contraction cracking, etc.) have been quantitatively evaluated. Geographic patterns of changes in the frozen soil layers as a factor of landscape and geological properties have been relatively well studied. Several models that attempt to forecast changes in the cryolithozone within different time frames during the 21 century have been developed.

Current assessments of the cryolithozone response to the contemporaneous and

mailto:vkonish@mail.ru

projected changes in climate do not

sufficiently account for heat exchange

between permanently frozen soils (PFS)

and the external environment. All external

factors, including climatic, do not impact PFS

directly as in the case on a glacial surface, but

act through a system of different layers, such

as snow, vegetation, top soil, active soil layer,

i. е., through landscape and its components.

Some authors differentiate factors of heat

exchange between PFS and the atmosphere

into temperature and non-temperature or

climatic and non-climatic.

Surface conditions and the intensity of

their influence change with annual seasons.

Directional climate change parameters

(specifically, an increase in the average annual

summer or winter air temperatures) influence

other components of the environment

affecting the heat exchange between

PFS and the atmosphere. Therefore, PFS

characteristics react to the entire complex of

parameters that are changing in response to

changes in the landscape climate.

An array of positive and negative feedbacks

appears as a result. In certain conditions,

these multidirectional feedbacks manifest

themselves in a varying intensity of the PFS

reaction to changes in air temperatures and

in a different character of the responses.

Numerous mechanisms that determine

these relationships are poorly understood at

the present time. For example, the influence

of snow cover on permafrost temperatures

is relatively well studied, but the influence

of the snow cover structure, its density,

porosity, and thermal-physical characteristics

that undergo winter changes and ultimately

influence the PFS thermal properties are

completely lacking in understanding.

Changes in surface conditions that

accompany warming or cooling may strongly

transform the direction of the freeze and

thaw process and the growth or degradation

of permafrost. In some landscapes, these

changes will act in the same direction as

the direction of climate trend enhancing its

influence, while in other conditions changes

will act in opposing directions resulting in a

weakening of the influence of climate trend

[Koreisha, Vtyurin, Vtyurina, 1997].

Background data on this important theoretical

conclusion have been obtained as early as in

1930th. A detailed analysis of spatial changes

in discontinuous cryolithozone in the basin

of the Selemdja river (Russian Far East)

established that "enhancement of permafrost

along with its degradation within the same

limited in size piece of land" [Kudryavtsev,

1939; Shvetsov, 1959] may be occurring

depending on geological, geomorphological,

hydrogeological, hydrological, and

geobotanical environment.

It may be assumed that spatial patterns

are analogous to temporal patterns of the

cryolithozone formation.

Despite the long established by permafrost

science fact that complex and ambiguous

dependencies between climatic and

permafrost characteristics exist, modern

literature still has elements of a simplified

approach to the analysis of the relationships

between thermal and other characteristics

of PFS, e. g., temperature and climatic

parameters.

Most of forecast models that describe

relationships between climate and permafrost

are mono-factorial and they only account for

direct links of the cryolithozone to isolated

environmental parameters, in particular, to

air temperature; some models, in addition,

include precipitation and thickness of the

snow cover at best.

The IPCC report [2008] contains examples of

such simplified approach to the assessment

of changes in the cryolithozone due to

global warming of the recent decades.

The purpose of the paper presented

herein is to demonstrate, using specific

examples from different regions of the

cryolithozone, ambiguous responses of

permafrost properties to two different-scale

climate warming events, i. е., warming of

recent decades and warming of the late

Pleistocene - Holocene.

CHANGES IN CRYOGENIC CONDITIONS

IN ISLAND AND MASSIVE-ISLAND

PERMAFROST ZONES

In the southern belt of the cryolothozone

in northern Europe, West Siberia, Canada,

and Alaska, permafrost may be found in

isolated or relatively extended areas of peat

hummocks and raised peat mounds that are

referred to as "palsas" in literature.

Recently, in such areas of different

countries (USA, Canada, Sweden, Finland,

and Mongolia), detailed studies on palsa

dynamics were conducted.

In 1933, a special field expedition was

organized by the Commission of the Academy

of Sciences of the USSR to identify the southern

permafrost border near the city of Mezen,

Kanin coastline of Mezen Bay. This expedition

obtained fundamental results on the dynamics

of frozen peat mounds [Datsky, 1937]

In this area, permafrost is represented by

1 -2 m high peat mounds 15 x 20 m in area.

Authors explain negative temperatures in

the peat mounds by the winter wind snow

removal off the tops of the hummocks. Some

hummocks are in a stationary state, but most

of the hummocks are degrading, i. е., they are

decreasing in size as a result of permafrost

thawing both on the slopes adjacent to

waterlogged lowlands and from the bottom.

However, the upper permafrost boundary

within this hummocky area, i. е., the thickness

of the STL, is relatively consistent (about

0,5 m) and exists at a permanently stable

state till the moment of almost complete

permafrost thawing and hummocks

degradation and collapse.

Therefore, hydrological properties, i. е., the

inundation level of wetlands that surround

the hummocks, not the temperatures,

may be the main factors that influence the

collapse of the hummocks.

The authors of this fundamental research have

still regarded the 1916-1930 climatic change

with temperatures 10 С higher than in 1883-

1915, to be the main factor of permafrost

degradation. In this work, for the first time a

connection between climate warming and the

degradation of permafrost in peat hummocks

has been established. Also, the research

has demonstrated that human impact (a

disturbance of hummocks surface, destruction

of vegetation, livestock grazing) may enhance

degradation processes in most cases.

Numerous research efforts that use modern

methods (time sequence aerial photography

comparison, geophysics, stationary

observations) to study peat areas and frost

mounds within the spread of island and

massive-island permafrost near its southern

boundary have been undertaken in different

areas of Finland, Sweden, Norway, Alaska,

and Canada. These studies have confirmed,

in general, the results of the field expedition

to Mezen region, and allowed to better

understand the climate change dynamics of

this type of permafrost. In all these regions,

a decrease in the area of frozen peatlands

represents a prevailing tendency that

reflects permafrost degradation. However,

the degradation of frozen peatlands occurs

in a lateral direction, i. е., from the sides,

and the speed of thermokarst development

depends upon the level of inundation of

surrounding wetlands. On average, the speed

of the degradation during the last 50 years

was 1% of the total area considered. Within

relatively well drained wetland sites, heaving

and growth of new peatlands and mounds

is occurring. Sometimes the degradation

of permafrost on one side is accompanied

by freezing, heaving, and the increase in

the size of the peat mass on the other side

[Kershow, 2003] Positive temperature change

trends do not practically influence the depth

of seasonal thawing within degrading or

growing peatlands. Some authors state that

the scale of permafrost degradation within

peatland areas does not at all depend on

changes in average annual temperatures,

but is influenced by local factors [Beilman,

Robinson, 2003].

Furthermore, certain observations prove that

permafrost dynamics within peat areas is not

defined by climate trend, but is determined

by a combination of specific summer and

winter weather patterns, which promote

permafrost preservation and the formation

of new frozen soils. The southern belt of the

cryolithozone of Timan-Pechora region during

the last 60 years [Osadchyi, Osadchaya, 2008]

may represent one of the examples. Another

example is the modern conditions of the

Lapland region in Finland, where the increase

in wind activity accompanied by the removal

of snow from the mounds 90 cm and higher

therefore promoting the formation of frozen

islands has led to the formation of new palsas

[Ronkvo, Seppala, 2003].

REACTION OF PERMAFROST TO MODERN

CLIMATE CHANGE IN OTHER REGIONS

Scientific data summarized by the IPCC

[2008] report an increase in permafrost

temperatures at the depth of zero annual

amplitude (tav) during recent decades in

different regions of the cryolithozone of

Alaska, northern Canada, northern European

part of Russia, northern Scandinavia, Tibetan

Plateau, Mongolia, and Tien Shan. The

increase in annual air temperatures along

with changes in the snow cover have been

suggested as the main factors of such

permafrost response.

These data describe permafrost developed

on mineral soils and characterize frozen

soils within autonomous landscapes; these

data differ from the results presented in the

previous section that pertain to frozen peat

areas.

In cases when temperature observations

in boreholes encompass a more complete

spectrum of landscape conditions of a site,

it appears that PFS respond to changes in

air temperatures (to warming) with varying

intensity and in a different pattern.

Earlier, we have demonstrated that positive

responses, i. е., positive feedbacks, to climate

warming, are present in autonomous or close

to autonomous, environments. However,

monitoring of permafrost temperatures in

Central Yakutia [Skryabin,Skachkov,Varlamov,

1999] indicated negative feedbacks in

subaqueous or superaqueous subordinate

environments, such as floodplains, low

terraces, shallow valleys, and inter-hummock

depressions [Konishchev, 2003].

Recently (the last 10 to 20 years), in the area

of Chulman basin (Aldan shield), geothermal

data showed a temperature increase tav of

0,1-0,3°C at the depth of 40 m in 23% of

the watershed boreholes. In most boreholes

located in intact natural conditions, a quasi-

stationary thermal regime is preserved at

the depths of 20-50 m, which does not

respond to the increase in average annual

temperatures during the last 20 years.

In some cases, in river floodplains and valleys

within sites with a pit-and-mound relief,

temperature decreased by 0,1 °C temperature

decrease (tav) to the depths of 20-40 m

[Zheleznyak, Zavadsky, Mitin, 2006] These

data undoubtedly confirm an existence

of an inter-landscape differentiation of tav

feedbacks to climate change.

Subordinate (subaqueous) landscapes may

have different feedbacks to climate warming

because, unlike autonomous landscapes,they

may develop protective responses including

an increase in the topsoil water content,

more intensive growth of hygrophilous

plants (moss), accrual of organic material in

the topsoil accumulative horizon, etc.

The extent of this phenomenon in relation

to different temperature and cryolithological

regions of the cryolithozone is hard to

evaluate at the present time due to a limited

volume of actual research data. In any case,

in the zone of discontinuous permafrost

(on mineral substrates), the landscape

differentiation in permafrost response

to climate warming and especially of its

southern part should be very specific.

In discontinuous permafrost area, a

temperature increase and melting of PFS

from the surface to the depth of several

meters, results in the development of

the residual thaw layer. This process was

detected by thermometric observations

in frozen mineral soils in the Urengoi gas

condensate field (West Siberia) [Vasiliev,

Drozdov, Moskalenko, 2008]. Unlike island

and massive-island permafrost where

frozen peat hummocks are degrading in

a lateral direction, in the northern part of

discontinuous cryolithozone, mineral layers

are degrading in a frontal direction, i. е., from

the top to the bottom.

Indicators of the cryolithozone feedbacks

to the modern climatic changes based on

monitoring include variations in the depth

of the seasonal freezing-thawing and the

temperatures of frozen soils at the depth

of zero annual amplitude (tav) (as a rule, 10

to 15 m). It is important to identify spatial-

temporal relationships between these

parameters.

Long-term stationary observations (over

30 ears in duration) in the northern part of

Western Siberia have showed that the depth

of the STL has a relatively weak response

to climate warming compared to annual

average temperatures of frozen soils (tav). The

temperature response of frozen soils within

some landscapes may be more pronounced

than a trend in air temperatures, which may

be influenced by an increase in multi-annual

snow deposits [Pavlov, Moskalenko, 2001].

Even more radical conclusion has been

reached in research activities conducted in

the northern part of West Siberia [Vasiliev,

Drozdov, Moskalenko, 2008]: a response of

dominant permafrost landscapes (peatlands

and wetlands) to climate warming may

be proceeding in two opposite directions,

i. e. maximal increase in seasonal thawing

accompanied by a minimal change in tav or

maximal increase in tav at a minimal change

in the STL thickness.

to 2007) the top soil and permafrost

temperatures at the depth of 0,7 m at the

site near Cape Franklin in Alaska increased

notably by 1,774°C and 2,34°C, respectively

following 1,334°C increase in air temperatures.

At the same time, the depth of the STL (with

an average thickness of 0,577 m) decreased

by 0,036 m. One of the main reasons for this

"discrepancy" is a growth of vegetation at the

medallion spots, or so called "greening effect

of the Arctic" [Daanen, Romanovsky, Walker,

and LaDouceur, 2008], which substantially

decreases absorbed solar radiation thus

decreasing the thawing index.

These examples indicate that the analysis

of the contemporary dynamics of the

cryolithozone under climate change

(warming) and forecast scenarios for

changes in the cryolithozone should

account for an entire array of properties of

a landscape and its isolated components

that change in response to the climate

change. The effects that counteract

manifestations of a leading process deserve

special attention. In different regions of

the cryolithozone, these relationships

differ and are very poorly understood;

their analysis should be based on regional

characteristics of the system "climate-

landscape-cryolithozone".

DYNAMICS OF PERMAFROST AT THE

END OF LATE PLEISTOCENE THROUGH

HOLOCENE

In northern Eurasia, a major scale degradation

of the late Pleistocene cryolithozone began

at the end of the late Pleistocene and the

Holocene interval. During the last 10 to

12 thousand years, the cryolithozne area has

shrunk by more than a factor of two. On the

Russian Plain and in West Siberia, the southern

boundary of the cryolithozone moved from

48° northern latitudes to its contemporary

boundary, i. е., several hundreds kilometers

to the north [Velichko, 2002].

Spotted (medallion) tundra permafrost has

a very specific response to the modern

warming. During the last 30 years (1977

Not least radical changes took place within

the modern area of the cryolithozone, the

most impressive of which is the evolution

Figure 1. Outcroppings of the ice complex

(low course of the Yana river, left bank, photographed by the author, 1967)

of the ice complex, i. е., ice-rich frozen soils

several dozens meters in thickness (Figure 1).

The ice complex sediments have most

widely spread in the late cryochron (15 to

18 thousand years ago), when sea regression

exposed the enormous areas of the Arctic

Ocean shelf zone, including the Laptev and

other seas. In these exposed areas and within

the territory of East Siberia, the accumulation

of the ice complex sediments took place. The

distribution of the ice complex sediments is

presented in Figure 2; the sediments of

the shelf zone areas have been destroyed

during the Holocene sea transgression and

are not shown on the map.The ice complex

sediments were forming under very severe

climatic conditions of the late cryochron;

in the northern part of Yakutia, average

temperatures of frozen soils (tav) were -25 ~ -28°C, reaching, in some areas, -30°C.

In Central Yakutia, the ice complex sediments

were forming at temperatures below

-10°C; with even lower air temperatures

[Konishchev, 1997]. At the present time,

the annual temperatures of permafrost in

these regions are -12 ~ -14°C and -3 ~ -5°С,

respectively.

A sharp increase in permafrost temperatures

was not the single consequence of the

cardinal climate change at the end of the late

Pleistocene cryochron - Holocene transition.

One of the main processes of the relief-

forming ice complex sediment evolution

during the last 12 to 13 thousand years

was their thermokarst erosion differentiation

and a formation of numerous thermokarst-

erosion depressions.

Currently, in the North and Central Yakutian

lowlands, the ice complex remnants (i. е.,

ostanetsy or yedomas) and alas basins several

dozens of kilometers in size represent the

dominant relief forms along with river valleys.

Published data contain various assessments

of areas of the ice complex remnants in North

and Central Yakutian lowlands,and alas basins

(thermokarst erosion). In the coastal lowlands

of Yakutia, thermokarst-erosion basins cover

approximately 60-70 % of the total lowland

area; in Central Yakutia, alase basins cover

10-20 % of the total region area [Bosikov,

1978]. Thermokarst-erosion basins in Yana-

Omoloi watershed area occupy 35-40 to

60-70 % of the lowland area [Gordeev, 1971].

Figure 2. Spatial distribution of ice complex deposits in Russia.

Compiled by the author and N.A. Koroleva based on [15, 28, 31 ]

In Yana-lndigirka lowland, thermokarst-lake surfaces are spread over 65-75 % of the area, whereas, remnants of the ice complex occupy only 25 % [Lomachenkov, 1965]. In the Kolyma lowland, remnants of the ice complex are found on approximately 4 1 % of the total lowland area. Alases in Central Yakutia (theTungulu terrace and its juncture with the Abalakh terrace) cover from 15-20% to 30-50 % of the ice complex area [Ivanov, 1981;Torgovkin, 1988].

Despite an approximate nature of these

assessments, their comparison allows one to

conclude that surfaces composed of the ice

complex in the lowlands of Northern Yakutia

are notably more affected by thermokart

processes than within the territory of the

Central Yakutia lowland.

Paleogeographic data on the Holocene history together with direct observations indicate that alas depressions are practically not developing at the present time in

Central Yakutia with the exception of alases that formed as a result of anthropogenic impact [Katasonov, 1979]. Alases actively grew due to the thermal abrasion in the past, specifically during a period of inundation of alas depressions in the late 1930s.

At the present time, an expansion of alas

basins is detected within the territory of

Yana-lndigirka lowland [Tolstov, 1966].

The main cause for the formation and development of alases in Northern Yakutia and further to the south of this area within the Central Yakutian lowland was not the warming or some other temperature factors alone, but the inundation of the territory expressed by the formation of lakes. If in the northern part, this factor was acting constantly from the very end of the late Pleistocene through the entire Holocene, the aridization period in Central Yakutia started in the first half of the Holocene and is continuing through the present time [Katasonov, 1979].

The accumulation of the ice complex

sediments has reached its completion at the

end of the late cryochron; from that time on,

apical surfaces of ostantsy (or yedomas) have

not been subject to denudation, including

thermokarst and erosion. This statement

is supported by large volume of data on

radiocarbon dating that indicate that the

completion of the ice complex sediments

accumulation took place at the very end of

the late Pleistocene [Ivanov, 1981, Kaplina,

1981].

The materials presented above support

the statement that during Holocene, the

transformation of the ice complex sediments

was affected by multidirectional processes

[Shur, 1984].

Within the entire territory, thawing of the ice

complex was occurring not in the frontal, but

in the horizontal direction (laterally) when

masses of the ice complex were subject to

thermal abrasion of lakes and to thermal

denudation from the sides.

During the inundation period at the end of

the late Pleistocene and in the Holocene,

conditions of apical surfaces of the ice

complex remnants made possible the

preservation and the enhancement of their

stability.

As early as in 1940 [Efimov, Grave, 1940], the

idea emerged that a protection layer exists

which is located below the STL and above

the layer of buried ice or, according to the

modern theory, above the surface of the

Pleistocene singenetic polygonal-venous

ice. Under a thick forest in inter-alas surfaces

of Central Yakutia, this layer exists in a

permanently frozen state and does not thaw.

Soon after forest removal, the thickness of

the STL increases by 30-40 % (at an average

initial value of 1,3-1,4 m) and the protective

layer begins to thaw. The thickness of the

protective layer is 1,5 to 2 m.

This highly important conclusion has been

fully supported by further observations. In

the coastal lowlands, the protective layer is

also broadly developed on the remnants of

the ice complex. Some authors call this layer

the "cover layer" [Kaplina, 1981 ], while others

refer to this layer as to the "transient" layer

[Shur, 1984]

This layer is characterized by a very high ice

content (up to 60-70 %).The ice is represented

by ataxic and reticulate cryostructures. The

layer is composed of aleurolites and, in most

cases, 1,5 to 1,7 m deep.

Literature sources explain the genesis of

this layer from the two points of view. One

theory suggests that the cover layer is a

relic of the Holocene climatic maximum

when the depth of the seasonal thawing

in the ice complex remnants has increased

to approximately 2 m. Later, during cooling

in the late Holocene, part of this layer has

frozen from the bottom bringing the depth

of the STL to its current values, i. е., 40-50 cm

[Kaplina, 1981, 18].

Other authors conclude that during the

Holocene climatic optimum, such increase

in the depth of the STL was not possible

because a beginning of the inundation

process and the development of wetlands

in the coastal lowlands in the Holocene not

only did not coincide with the beginning of

warming, but had occurred in advance of the

warming. Indeed, irrefutable evidence exists

now that mass development of thermokarst

has begun not in the Holocene climatic

optimum, but substantially earlier, in the late

Pleistocene [Romanovskii, Gavriloy, Tumskoy,

et al., 1999]. In the remnants of the ice

complex (yedomas) hydrophilic vegetation,

including mosses and small shrubs,

expanded; peat-gleic soils have formed and

the STL soil water content has increased; all

these processes by no means promoted the

growth of the STL thickness.

In isolated and relatively rare cases, the

thickness of the cover layer in the ice complex

sequence somewhat exceeds average

values; at some sites, the cover layer forms

2,5-3 m deep moulds. This process was also

described by V.V. Kunitsky [2007] who used

the term "bylar" to describe ice deposits that lie below the STL and associated their formation with the STL of higher water content compared to background values. A talik does not form in this case. Yu.L. Shur [1988] indicates that during Holocene, in places where inundation has not reached the extreme levels expressed by waterlogging of the territory and lakes formation, hydrophilic vegetation have developed promoting permafrost preservation. The inundation levels in the ice complex remnants varied; within small depressions, inundation level was somewhat higher leading to greater thickness of the cover layer.

In the past, for some period of time, a background depth of the STL in surfaces composed by the ice complex was greater than in the modern time. During the late Pleistocene, the Sartansk glacial sediment layers of the ice complex have formed. This theory was suggested by G.F. Gravis [1969] and further developed by Yu.L. Shur [1988] who based his analysis on the spore and pollen data of the ice complex sediments that indicate the prevalence of grassy vegetation and greater than modern depth of the seasonal thawing in open landscapes of the steppe-tundra.

Recent research activities confirm this point of view. Bio-geographic data on fossil insect associations in the ice complex sediments indicate the presence of multiple species of steppe beetles with strict requirements to the heat supply, in addition to boreal species of the northern tundra. During sea regression that has spread in the late glacial maximum to several hundreds of kilometers to the north of the modern shoreline, within exposed areas, continentality increase resulted in higher summer temperatures compared to the temperatures of the modern tundra. In the coastal lowlands to the west of the Kolyma river, July temperatures were 5-7°C higher and reached 11-12°C [Alfimov, Berman, 2004].

During the regression of the Arctic seas, the

polar day effect (i. e„ the thawing index equal

to the sum of summer air temperatures) increased substantially due to a substantial cloud cover decrease (at the present time, the cloud cover is 65-70%) which enhanced the thermal impact of direct solar radiation.

As a result, the thickness of the STL grew substantially with a significant decrease in the average annual soil temperatures due to a strong cooling effect of polar nights when air temperatures were dropping to -70°C. Therefore, summer temperatures did not prevent but, on the contrary, promoted the spread of thermophilic ecosystems beyond the boundaries of the modern land area of the shelf zone providing sustenance (food) for mammoth fauna.

Burrows of fossil rodents may serve as another indicator of the depth of the STL during accumulation of the ice complex sediments [Gubin, Zanina, Maksimovich, Kuzmina, Zazhigin, 2003]. The thickness of this layer during accumulation of the Karginsky ice complex sediments is 80-85 cm (sedimentary outcroppings of Duvannyi Yar, lower course of the Kolyma river) and 60-65 cm (sedimentary outcroppings at Zelenyi Mys, lower course of the Kolyma river). During Sartansk period, when continentality reached its maximum, the thawing index also reached its maximal values and the thickness of the STL has increased to at least 1 m.

After the completion of the accumulation of the sediments of the ice complexes and during its further presence in the Holocene time and as a result of changes in landscape characteristics (an increase in moisture content, development of moss cover, etc.) within non-thawed sites, the STL depths decreased followed by the rise of the upper surface of permafrost and formation of the ice-rich transient layer [Shur, 1988]. The process of formation of this transient layer is similar to a process, known as a "quasi-syngenesis process" in the literature [Kudryavtsev, 1978; Shur, 1988].

However, this process is not the only

process that determines the formation of

Figure 3. Schematic representation of the transient layer formation during cyclic freezing-thawing.

Laboratory experiments with kaolin clay [8]; initial moisture content - 38%, mass volume - 1,3 g/см3,

I - V - cryogenic structures (black color) during freezing-thawing cycles.

v T H W , v sif - position of the thawing front and segregation ice formation

the cryogenic structure and the size of the

intermediate layer.

Water migration from the STL to the underly

ing upper horizon of permafrost is also of a

major significance. Water migration occurs

during summer seasons when the active

layer frozen during winter months, thaws and

a temperature gradient between its thawed

part and still frozen underlying permafrost

appears. This process is relatively well stud

ied in laboratory experiments and in field

observations [Parmuzina, 1978; Ershov, 1979,

Konstantinov, 1991, Solomatin, 1994]. The

field studies demonstrate that permafrost

water content may increase by a factor of

two reaching 50 % and higher as a result of

the downward migration of water from the

STL to the permafrost upper horizon. N.N. Ro

manovsky [1993] indicates that the migration

from the STL to the underlying frozen soils is

occurring within the range of average annual

temperatures (tov) of the frozen soils is equal

to -2 ~ -3°C to -8 ~ -10°C At tav close to 0°C

and due to small gradients, and at tav below

-10 ~ -12°C due to a low unfrozen water con

tent, this process is substantially weakened.

The significance of water migration process

to the transient layer that underlies the STL

becomes apparent from the comparison of

the data on numerous boreholes and out

croppings of the transient layer (1,5-1,7 m

on average) and the depths of the STL during

final phases of the ice complex accumulation

(approximately 1,0 m). Water was migrating

not only to the transient layer (which thick-

ness does not exceed 0,4-0,6 m), but to the

lower horizons as well.

The cryolithological composition of the ice-

rich transient layer of dark-gray or blue-gray

color significantly differs from the underlying

Sartansk horizon of the ice complex brown

in color and with massive and micro- fine

schliere cryogenic texture.

The frozen soil temperatures (tav) during

the formation of the layer composed of

brown aleurilites was very low, from -22 to

-33°C; summer temperatures of the upper

permafrost horizons did not exceed 10°C

[Konishchev, 2002]. This condition prohib

ited the migration of water from the STL to

the underlying frozen soils. Water migration

became possible either during the Holocene

or during one of the phases of warming at

the very end of the Pleistocene (Alleroed,

Boiling) when average annual and sum

mer temperatures increased sharply (tav =

= -8 ~ -14°C). Observations in other region

of the cryolithozone demonstrate that at the

present time, water is also migrating form

the STL to the underlying horizons of the

frozen soils [Chizov, Derevyagin, Mayer, 2005;

Konstantinov, 1991].

The formation of the transient layer as a

result of a transition of a part of the STL to

the permanently frozen state is, therefore,

complicated by water migration from the

STL to the underlying frozen layer.The overall

result is the increase in the ice content at the

surface of the remnants of the ice complex.

This process, together with the decrease in

the depth of the STL during the transition

from the upper Pleistocene to the Holocene,

may be interpreted as the aggradational

process that took place in parallel to the

degradational processes, i. e. an increase in

the area of alas depressions due to the ice

complex thermal abrasion, thawing of the

frozen horizons under lakes, etc. The ice-

rich transient or the "cover" layer is strongly

developed on the surface of the ice complex

and contemporaneously plays an important

protective role to external, first of all, thermal

impacts resulting from climate warming.

In the context of this analysis, the ice-rich

horizon of the upper permafrost horizon

could logically be called not a cover or

transient layer, but a protective layer, i. е., the

term that was initially suggested [Efimov,

Grave, 1940] could be used.

Thus, the destructive role of the Holocene

warming and the thermokarst erosion

modification of the ice complex sediments

represent only one type of reaction of the

cryolithozone to climate warming.The other

type of reaction are stabilizing changes in

landscape characteristics, e. g. the increase in

the moisture content of soils, development of

hydrophilic moss vegetation, accumulation

of organic matter in top soils, etc.

The presence of the protective layer provided

for the preservation of a significant part

of the ice complex during for more that

10 thousand years and it, in many ways,

determined modern landscape-permafrost

and geo-ecological conditions within a

large territory of East Siberia. Without this

protective layer, thermal abrasive and

thermal denudation processes could have

been many fold more intense, which can

be currently observed along the shoreline

of the Laptev and East Siberian seas where

the sea-banks composed of the ice complex

are intensively destructed by the sea thermal

abrasion and are retreating at a speed of

more that 10 cm per year.

At the present time, alases of Central Yakutia

are developing very slowly. The leading role

in the development of the thermokarst relief

belongs to anthropogenic activities [Bosikov,

1978].

Reliable data on somewhat significant

thermokarst processes in natural, intact by

human activities, ecosystems within inter

alas surfaces of Central Yakutia and the

apical surfaces of the ice complex remnants

in the Primorsky and the Northern Yakutia

resulting from the contemporaneous

climate change are absent. When surface

conditions are disturbed due to forest fires,

transport, forest logging, or agricultural

development in the area and within

surfaces composed of the ice complex

sediments, the depth of the STL increases

together with sagging, linear erosion and

other destructive processes that indicate

degradation of the permafrost complex.

Such reports often appear in the literature

[Gavriliev, Ugarov, 2008].

CONCLUSION

The response of the most important

characteristics of the cryolithozone to climate

warming depends upon the entire complex

of parameters of changing landscape and its

components that follow climate change.

The cryolithozone climate parameters

do not always directly correspond to the

parameters of atmospheric climate. The "soil-

ground" climate is determined by landscape

conditions where atmospheric climate

represents only one of the components.

Within some types of the permafrost

landscapes under climate warming,

protective reactions develop as a result

of negative feedbacks that stabilize some

permafrost characteristics (for example,

the depth of the seasonal thawing) or

even lead to a decrease in average annual

temperatures of the permafrost layer (tav), increase in the ice content of the upper

horizons of permafrost, and formation of

the ice-rich protective layer that mitigate

both inter-annual and multi-annual

variations in the depth of the STL. Under

climate warming, permafrost aggradational

changes (a decrease or stabilization of the

depth of the STL, decrease or stabilization of

t av in subordinate landscapes, and increase in ice content of the upper permafrost

horizons) are observed together with

degradation processes (increase in tav in autonomous landscapes, increase in the

area of taliks, etc.)

A response of permafrost to climate

change differs in principal within different

temperature and cryolithological types of

the cryolithozone, i. е., within the spread

of island and massive-island permafrost

represented predominantly by peatlands

on one hand, and within the spread of

discontinues permafrost where inter-fluvial

landscapes composed of mineral soils

(autonomous) alternate with waterlogged

subordinate landscapes of depressions

and river valleys (floodplains), on the other

hand.

The ambiguous response of permafrost

characteristics to climate warming is

observed both within the context of the

recent decades and of the longer time

intervals, specifically, during the period of

the upper Pleistocene - Holocene.

The integrity of the ice complex sediments

very unstable to thermal impacts, represents

a typical relict of the late Pleistocene cryochron

and is a consequence of the impact of

several permafrost processes that have led

to the formation of the protective layer at

the ice complex remnants in response to the

Holocene warming.

The discussion presented above supports

the need for further research based on

detailed monitoring in different regions

of the cryolithozone of Russia and other

countries. •

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Konischev, Vyacheslav Nikolaevich, was born in Moscow, in 1938. In 1960, he graduated from the Department of Polar Countries (renamed in 1967 to the Department of Cryolithology and Glaciology) of the Geography Faculty of M.V. Lomonosov Moscow State University. He received the Doctor of Science in Geography degree in 1978 and became a full Professor in 1981. He was awarded the title of Honored Worker of Science of the Russian Federation in 1999.

Professor Konischev has served as the Chair of the Department of Cryolithology and Glaciology since 1993. The areas of his scientific and teaching activities include permafrost science, cryolithology (composition and structure of cryolithic zone), and history and geo-ecology of Earth cryosphere. He has participated in scientific field expeditions in different regions of Northern Eurasia and Northern America; delivered lectures in universities of Canada, Iceland, Poland, China, and Japan; and published nine monographs and textbooks, and over 200 papers in peer-reviewed journals and proceedings.