Results & Discussion

Climatic changes in Altai Mountains

Seasonal changes of temperature and precipitation in Altai Mountains are shown in Figure 10. High temperatures and a large amount of precipitation are typical in summer (June, July, August). Spring temperatures are slightly higher than in autumn, and precipitation is observed more in spring (Mart, April, May) than in autumn (September, October, November). The lowest temperatures are recorded in winter (December, January, February), as is the minimum amount of precipitation.

Figure 10. Average monthly temperature and precipitation in Altai Mountains

Change in temperature

The Mann-Kendall statistical test showed that there is a significant increasing trend (0.43 ⁰C/10yr) in the mean annual temperature of the Altai Mountains during the 1957–2017 (Table 2). Significant warming is observed in spring (0.64⁰C/10yr) and autumn (0.37⁰C/10yr), which is consistent with the results for Asia in the Fifth Assessment Report [IPCC, 2014; Hijioka et al., 2014]. The warmest year was in 1965 (0.47⁰C), and the coldest in 1969 (-2.84⁰C), as well as the coldest winter (-21.95⁰C), while the warmest (-12.78⁰C) was in 1988. Large differences between the coldest (-3.37⁰C in 1996) and the warmest (0.98⁰C in 1981) spring is not observed, and there were big differences for the autumn (-4.58⁰C in 1974 and 1.12⁰C in 2006). Differences in temperature for the summer amounted to just over 2⁰C (13.14⁰C in 1971 and 15.72⁰C in 1965).

Table 2. Annual and seasonal temperatures (⁰C) and corresponding climate tendencies in Altai Mountains (1957-2017)

  mean min max climate tendencies

(⁰C/10yr)

Annual -0.84 -2.84 0.47 0.43
DJF -15.76 -21.95 -12.78 0.25
MAM -1.03 -3.37 0.91 0.64
JJA 14.84 13.14 15.72 0.21
SON -1.42 -4.58 1.12 0.37

The study of temperature maps made according to NCEP/NCAR reanalysis data showed that the coldest temperatures are observed in the central parts of Altai (Figure 11). In winter, the cold center shifts to the Mongolian part (Figure 12). In the spring again, the coldest temperatures are typical for the central part, as for the annual as a whole (Figure 13). In the summer, the center of the cold moves to the Russian part (Figure 14), and in autumn it has two main parts in the Russian and Mongolian Altai (Figure 15).

Figure 11. Map of temperature annual (1957-2017)

Figure 12. Map of temperature winter (1957-2017)

Figure 13. Map of temperature spring (1957-2017)

Figure 14. Map of temperature summer (1957-2017)

Figure 15. Map of temperature autumn (1957-2017)

In this project we studied two periods: the first period from 1957 to 2012, i.e. the closing date as in the Fifth Assessment Report [Hijioka et al., 2014]; the second period from 1957 to 2017, that is, the interval when there are meteorological data. The Mann–Kendall–Sneyers test has found in addition the dates of the step change points of temperature. Figure 16 illustrates that in the first (1957-2012) period this occurred around 1980, while in the second (1957-2017) period around 1996 (Figure 17). The results for the first period (1957-2012), namely, the increase in warming around the 1980s, are in good agreement with the results obtained for Asia [Hijioka et al., 2014]. In the analysis of a longer period (1957-2017), the point of temperature gradient shifted by almost 15 years (1996). That is, at the beginning of the last decade, more intensive warming began, than the previously released from 1980. Spatial analysis of temperature maps showed that in the first period (1957-2012) the field of the coldest temperatures was characteristic for the first interval (1957-1979) (Figure 18). However, in the second analysis period (1957-2017), the temperature field in the first interval (1957-1995) had two centers located in the Russian and Mongolian Altai, which were not determined in the second interval (1996-2017) (Figure 19). This result confirms that the second period was significantly warmer than the first period.

Figure 16. Annual temperature and mean during 1957-1979 and 1980-2012 (dotted lines) in Altai Mountains

Figure 17. Annual temperature and mean during 1957-1995 and 1996-2017 (dotted lines) in Altai Mountains

Figure 18. Map of temperature annual (1957-1979 – overhead; 1980-2012 – down)

Figure 19. Map of temperature annual (1957-1995 – overhead; 1996-2017 – down)

The Mann–Kendall–Sneyers test has not found the dates of the step change points of temperature in winter (Figure 20). While for spring, summer and autumn, the step change points were also found occurred in 1996 (Figure 21-23). The mean, minimum and maximum data for two intervals is in Table 3.

Figure 20. Winter temperature and mean during 1957-2017 (dotted line) in Altai Mountains

Figure 21. Spring temperature and mean during 1957-1995 and 1996-2017 (dotted lines) in Altai Mountains

Table 3. Annual and seasonal temperatures (⁰C) in Altai Mountains during 1957-1995 and 1996-2017

  mean min max
1957-1995 1996-2017 1957-1995 1996-2017 1957-1995 1996-2017
Annual -1.15 -0.29 -2.84 -1.88 0.47 0.73
DJF
MAM -1.44 -0.31 -3.34 -3.44 0.91 2.29
JJA 14.53 15.40 13.14 14.03 15.72 16.33
SON -1.88 -0.60 -4.58 -3.13 1.09 1.12

 

Figure 22. Summer temperature and mean during 1957-1995 and 1996-2017 (dotted lines) in Altai Mountains

Figure 23. Autumn temperature and mean during 1957-1995 and 1996-2017 (dotted lines) in Altai Mountains

The most significant increase in temperature when comparing two intervals (1957-1995 and 1996-2017) occurred in autumn (1.28⁰C) and spring (1.13⁰C), while for annual and summer the average temperature increase was approximately the same (0.86). A significant increase in minimum temperatures was obtained for the autumn (1.45⁰C), while the maximum for spring (1.38⁰C).

 

Change in precipitation

With the temperature increase in Altai Mountains, the mean annual precipitation shows increase, but various tendencies in different seasons (Table 4). Positive tendencies in summer (2.9 mm/10yr) and spring (1.4 mm/10yr), and negative in autumn (-0.7 mm/10yr) and winter (-0.1 mm/10yr). The maximum precipitation was in 1984 (597 mm), as in the summer (284 mm). Winter precipitation was the most in 1966 (96 mm), in the spring in 1995 (149 mm), and in the autumn of 1987 (158 mm). The minimum precipitation was in 1962 for the annual (358 mm), in 1967 for spring (67 mm), in 1974 for summer (127 mm), in 1997 for autumn (59 mm) and in 2005 for winter (33 mm).

Table 4. Annual and seasonal precipitation (mm) and corresponding climate tendencies in Altai Mountains (1957-2015)

  mean min max climate tendencies

(mm/10yr)

Annual 472 358 597 3.3
DJF 59 33 96 -0.1
MAM 114 67 149 1.4
JJA 192 127 284 2.9
SON 107 59 158 -0.7

The precipitation maps showed that that the fields with maximum annual precipitations are located in the north-west part Russian Altai (Figure 24). In winter and spring, the highest precipitation was recorded in the north of Altai Mountains (Figure 25-26), but the precipitation is also high in the west. The maximum precipitation in Russian part Altai is determined in the summer (Figure 27), a similar distribution is observed in the autumn (Figure 28).

Figure 24. Map of precipitation annual (1957-2015)

Figure 25. Map of precipitation winter (1957-2015)

Figure 26. Map of precipitation spring (1957-2015)

In this project, we studied two periods: the first period until 2012 and the second until 2015. For the first period (1957-2012), there were no step change points precipitation for annual or for the seasons. The Mann–Kendall–Sneyers test has found of the step change points of annual precipitation occurred in 1980 for the second period (Figure 29-30 ). The comparison of precipitation maps shows large parts covered by maximum precipitation in the second period (1980-2015), compared to the first period (1957-1979), which confirms the calculation of the Mann–Kendall–Sneyers test. However, for the winter, the step change points precipitation (Figure 31) were not distinguished, as for the temperature. In spring (Figure 32) and summer (Figure 33), step change points precipitation was detected as well as for the annual precipitation in 1980. For autumn, the trend of step change points precipitation was not revealed, as for winter (Figure 34). For the annual, spring and summer seasons the mean, maximum and minimum precipitation are in Table 5.

Figure 27. Map of precipitation summer (1957-2015)

Figure 28. Map of precipitation autumn (1957-2015)

Figure 29. Map of precipitation annual (1957-1979 – overhead; 1980-2015 – down)

Figure 30. Annual precipitation and mean during 1957-1979 and 1980-2015 (dotted lines) in Altai Mountains

Figure 31. Winter precipitation and mean during 1957-2015 (dotted line) in Altai Mountains

Figure 32. Spring precipitation and mean during 1957-1979 and 1980-2015 (dotted lines) in Altai Mountains

Figure 33. Summer precipitation and mean during 1957-1979 and 1980-2015 (dotted lines) in Altai Mountains

Figure 34. Autumn precipitation and mean during 1957-2015 (dotted line) in Altai Mountains

Table 5. Annual and seasonal precipitation (mm) in Altai Mountains during 1957-1979 and 1980-2015

  mean min max
1957-1979 1980-2015 1957-1979 1980-2015 1957-1979 1980-2015
annual 455 482 358 430 559 597
DJF
MAM 111 116 67 92 140 149
JJA 175 201 127 129 222 284
SON

The maximum increase in mean precipitation from the first to the second period was obtained for the annual and summer (more than 25 mm), a smaller increase for spring (5 mm). The increase in minimum precipitation for the annual was 72 mm, while for maximum precipitation in summer – 62 mm.

Climate changes in Western Mongolia

The analysis of the average annual temperature showed its steady positive growth in all weather stations-for the period from 1958 to 2017, it was 2.3 °C; the highest value of 3.1°C is observed at the station Hovd, located at medium-high levels in the Central part of the study area. The average annual temperature is negative and is -1.06°C for four weather stations in Western Mongolia. The average annual temperature is positive only at the weather station Ulgi. The standard deviation by year is about 1 °C for all the stations listed (Table 6, Figure 35).

Table 6. The changes of Temperature and Precipitation in the weather stations located on the territory of the Western Mongolia during the period from 1958-2017

 

The name of the weather station

 

Elevation, m ASL

 

Ratio of temperature trend

 

Temperature change during the study period, 0С

 

Average annual temperature, 0С

 

Standard deviation, 0С

ALTAI 2181 +0,037 +2,07 -1,01 0,96
ULIASTAI 1759 +0,026 +1,53 -1,85 1,01
ULGI 1715 +0,039 +2,22 +0,74 1,24
HOVD 1405 +0,053 +3,07 -0,96 1,37
ULAANGOM 939 +0,037 +2,03 -2,53 1,19

The highest values of average annual temperatures occur in the period 1990-2000, in which there is a fluctuation in the amplitudes of average annual temperatures from the mean annual value of 2 °C to 5 °C.

Table 7. The changes of Temperature and Precipitation in the weather stations located on the territory of the Western Mongolia during the period from 1984-2017.

 

The name of the weather station

 

Elevation, m ASL

 

Ratio of temperature trend

Temperature change during the study period, 0С
ALTAI 2181 +0,048 +1,54
ULIASTAI 1759 +0.044 +1,41
ULGI 1715 +0,057 +1,82
HOVD 1405 +0,059 +1,89
ULAANGOM 939 +0,050 +1,60
OMNO-GOBI 1590 +0,030 +1,29
BARUUNTURUUN 1232 +0,052 +1,66
BAITAG 1186 +0,048 +1,53

Figure 35. Average annual temperature (0С) dynamics for the period from 1958 to 2017 for five stations: 1. Altai, Hovd, 3. Ulaangom, 4. Ulgi, 5.Uliastai.

Table 8. Temperature and precipitation changes, in the weather stations located on the territory of the Western Mongolia during the period from 2000-2017.

The name

of the weather station

Elevation, m ASL Ratio of temperature trend Temperature change during the study period, 0С Average annual precipitation, mm Standard deviation, mm Variation coefficient,

%

ALTAI 2181 +0,017 +0,27 182,37 54,90 30
ULIASTAI 1759 +0,019 +0,30 210,49 62,66 30
ULGI 1715 +0,019 +0,30 112,85 25,19 22
HOVD 1405 +0,017 +0,27 124,00 46,43 37
ULAANGOM 939 +0,013 +0,21 133,34 42,70 32
OMNO-GOBI 1590 +0,021 +0,36 133,34 42,70 32
BARUUNTURUUN 1232 -0,013 -0,21 220,52 55,43 25
BAITAG 1186 +0,077 +1,23 90,29 41,35 46
ERDENI 2417 +0,163 +1,90 60,01 25,70 43
TOLBO 2101 +0,043 +0,67 187,78 112,43 60
TONHIL 2095 +0,053 +0,74 97,97 30,34 31
NOGOONNUUR 1480 +0,126 +0.04 98,58 33,17 34
URGAMAL 1263 +0,135 +1,89 99,18 37,25 38
HUNHATAOORTOO 1051 +0,083 +1,20 127,06 80,37 63

Figure 36. Average summer temperature dynamics (from June to September) for 1958-2017

The average annual temperature may drop to minus 5 °C in some years (e.g., Altai weather station), and in some years rise to plus 3 °C (e.g. weather station Hovd). Data for the period 1960-70 are incomplete for some weather stations, therefore have a discrete character. Analysis of data for the 30-year observation period also showed a positive trend with angular coefficients from 0.030 to 0.059; for this period, the temperature increase is less than for 60 years period — the average warming is 1.6 °C (Table 7).

The Weather data are presented for a larger network of weather stations, for the period from 2000 to 2017 (Figure 37, Table 8), which fairly evenly covers the territory of Western Mongolia. A number of data analyzed by the thermal regime showed a further steady increase in average annual temperatures, with the exception of the weather station Baruunturuun, located in the North-East of the study area. The positive growth of average annual temperatures is observed in all stations-about 0.6 °C, the highest value of 1.9 °C is the weather station Erdeni, which has the highest absolute altitude (2417 m).

The dynamics of the average temperature fo the ablation period (average temperature from June to September) and the dynamics of the amount of winter precipitation (the amount of precipitation from November to April) were analyzed additionally (Figure 36, 37).

 

The average annual air temperature growth was 2.6°C for weather stations Hovd for the 60-year period; for Omno-Gobi — 1,29°C, for 30-year period; for Tolbo — 0,67°C, for the last 18 years. The value of total winter precipitation is high oscillatory in the years of observations, both in the area of the investigated ridge and throughout Western Mongolia.

Figure 37. Dynamics of average summer temperatures and the amount of winter precipitation during the 2000 – 2017

 

Landscape Change in Permafrost Regions of Mongolian Altai

Geothermal conditions of the Mongolian Altai have spatial heterogeneity, which has caused a wide variety of cryomorphogenesis processes, the development of which is facilitated by low soil temperatures and the difference in the depth of the active layer.

Field geothermal observations (Figure 38) revealed that the minimum depth of the spring-summer thawing is fixed in peaty rocks is 0.8 m on the test site of Ehen-Nuur in 2017. The maximum depth of the active layer (up to 5 m) was observed in coarse-grained, boulder-pebble, limnic and fluvio-glacial deposits.

The depth of the active layer is in the range of 1-1.7 m in deluvial-proluvial sediments (rubble with loamy filler), in the morainic boulder-loams 1.5-2.5 m, and in the clay — about 1 m.

Figure 38 . Drilling of borehole for permafrost geothermal monitoring on Ehen-Nuur key location and thermal data from borehole logging system.

The new borehole up to 3.5 m deep was drilled on Ehen-Nuur key site at the 2106 m a.s.l., a thermistor chain and logging system were installed in the borehole. The thermistor chain installed in borehole consists of 35 negative temperature coefficient resistors type YSI 44031. Their accuracy without calibration is +/- 0.1C over the whole measurement range and  there is no significant drift at low temperatures. The sensors in the chain are installed in steps of 10 cm.

The analysis of thermal data from this borehole (Figure 38) showed that active near-surface freezing of the active layer began in October and reached maximum depths by the second decade of February, during the observation period of 2017. Stable thawing of seasonally frozen layer was observed in April. The top layer is about 20 cm deep warmed up to 15 °C to the middle of summer, but only in August this temperature penetrated to a depth of 80-85 cm. The maximum thawing up of ground to a depth of 2 m was recorded at the end of September: rocks at this depth warmed up to 0.5 °C.

A pronounced high-altitude differentiation is observed in the distribution of cryogenic landscapes within the studied ranges of Sutay and Tsambagarav. Forms of frost weathering: blockfields, blockstripes, altiplanation terrace are typical for relatively well-moist areas of slopes, within the subnival zone, where the temperature of the near surface layer of air and the upper horizons of soils often pass through 0 °C (Figure 39). Frosty sorting of stone material is occure on similar hypsometric levels, within the aligned or inclined (up to 4°) surfaces, with the formation of stone polygons and rings up to 2 m in diameter, formed frost scars. Stone polygons are marked by us on the watersheds of the South-Eastern slope of the Sutay ridge, folded from the surface with loose rubble sediments with inclusions of coarse-grained material.

Figure 39. The blockstripes on the left side of Khar-Asga valley (Massive Tsambagarav). A.s.l. 2750 m

Sharp fluctuations in the temperature of the surface air layer during the year lead to the appearance of discontinuous deformations in the surface layer of soil and frost crack formation, which is most actively occurring on the surface of river terraces, flat areas of slopes and bottoms of local depressions. The cracks that appear in this case have a depth of 2 meters, with a width in the upper part up to 15 cm. The processes of cryogenic heaving occur at similar with cracking geomorphological levels. Precipitation mostly by the end of summer supports the development of cryogenic heaving; such moisture contributes to a

significant moisture saturation of the surface layer of the soil by the beginning of the freezing season. Seasonal (up to 1 m high) and perennial mounds up to 2-3 m high are distinguished among various forms of cryogenic heaving.

Figure 40. Young stone rings at the bottom of dry thermokarst hollow. Namarzhaan valley, Massive Tsambagarav, a.s.l. 2520 m

The processes of thawing of ground ice from the day surface play an important role in the transformation of landscapes in the mountainous areas of the Mongolian Altai. In this regard, a special emphasis in the research was made on the assessment of the rate of development of thermokarst in modern conditions.

The dynamics of thermokarst processes is studied using the materials of polychronous space survey and ground observations. The analytical processing of remote sensing data showed a widespread and steady increase in the number and area of thermokarst lakes within the moraine complexes of the little ice age of the Tsambagarav massif, the processing performed in the environment of GIAS «EvCLiD» for the period from 1968 to 2016 (Figure 41, 42).

The tendency of shallowing and reduction of the water area of the old thermokarst lakes is noted at lower hypsometric levels, in zone of discontinuous permafrost, with the simultaneous appearance of young, due to intensive subsurface thawing of ice-bearing loose sediments (Figure 43, 44).

Figure 41. Comparative GIS analysis of Tumurt mountain-glacial basin (Tsambagarav massive). The blue line on the upper image (sensor KH-4B, 1968-08-11) showing position of Tumurt glacier terminus in 2015. The red lines on the lower image (sensor WorldView – 110, 2015-08-19) indicate the contours of lakes formed within LIA moraine from 1968 to 2015.

Figure 42. In the Altai mountains, and specifically in the Tsambagarav ridge a marked decrease of the glaciated area has occurred since the end of the Little Ice Age, and it has been accelerated since the last decades of the 20th century. As a result of the glacier retreat new thermokarst lakes (upper image) and pro-glacial lakes(lower image) are originated, and often the area and volume of existing ones increases.

Figure 43. Comparative GIS analysis of Ehen-Nuur key site. The left image (sensor KH-4B, 1962-08-22) shows the state of thermokarst lakes in 2016. The red fill on the right image (sensor WorldView – 110, 2015-08-19) indicate the contours of the lakes in 1962 and green-blue fill in 2016.

Figure 44. The photo, looking north, at what remained of the bay of Lake Ehen-Nuur on September, 2017. Prior to 1962, the lake occupied the entirety of the now-dry lake bed

Traces of glaciers advance in the Little Ice Age (17-19 centuries) expressed in the form of terminal and lateral moraine complexes clearly preserved in the relief of the valleys of the ranges Sutay and Tsambagarav. The main morphological feature of the complexes is the existance of a ridge of the frontal moraine, reliably recognized by the materials of space survey, and used by us as a reference point for the reconstruction of the spatial characteristics of the nival-glacial zone for the period of the Little Ice Age.

Figure 45. The Little Ice Age moraine below the right-side valley glacier in the Yamaat mountain-glacier basin, Tsambagarav massive, Aug 2017 (Upper photo) and frontal part of the Little Ice Age terminal moraine of Tsagangol glacier, Sutai ridge, Aug 2017 (Lower photo)

Glacial landscapes occupied 16.02 sq. km ridge Stay and were distributed on 99,104 square kilometers within the range Tsambagarav in the maximum transgressive stage of the little ice age, based on 3D topography. The decrease in regional temperatures by 0.6-0.8 °C caused the inversion of the lower boundary of the ice belt, which reached 200 meters. Thickness of valley glaciers in the basin of the river Tsagangol two times superior to the modern, in a number of mountain-glacial basins. The thickness of glaciers was reconstructed by the hypsometry of lateral moraines

Figure 45. The red line on the image show Little Ice Age glacial extension in the Sutai ridge. Blue colour filling show the position of glaciers in 2017.

The climate change in the post-maximum phase of the Little Ice Age resulted in a spatial transformation of the nival-glacial belt of the ridges, expressed in a progressive reduction in the size of the glaciation and the uplifting of its lower vertical limits. The total area of the Sutai ridge glaciation has decreased to 11.21 sq km, and the Tsambagarav massive to 66.57 sq km, by August 2017.

Subglacial deposits moved to the category of subaeral, in the deglaciation zone, under the influence of the changed climatic background. Embryonic periglacial landscapes began to form on a young lithogenic basis. The space-time dynamics strictly subordinated to the change of external hydrothermal conditions that is the main property of landscapes developing within the periglacial zone of the Sutai and Tsambagarav ridges. Physical weathering, soilfluction, cryogenic slumping and thermokarst, are among the most important processes involved in their formation.

Figure 47. The extends of glaciers of Tsambagarav massive in the  Little Ice Age maximum (Upper image) and position of glaciers in 2017 (Lower image)

Solifluction forms are spread widely, the study area, and confined to the lower parts of the slopes of the valleys in the altitude interval from 2400 to 3000 m. Their genesis and dynamics are associated with the wide spread of permafrost loose rocks, hydrothermal regime of the region and the development of vegetation. As a rule, solifluction forms occur in groups and occupy convex parts of the sides of valleys with angles from 10 to 30°. They have the form of festoons; larger forms are represented by terraces. The size of the terraces varies widely: length — from 4 to 30 m, width from 0.5 to 6 m and height of the scarp from 0.5 to 1.5-2 m. Open, not covered forms of solifluction formations are confined to higher hypsometric levels of slopes, often with a pronounced scarp of rough debris. At the base of the slopes, more widespread solifluction grass-covered terraces and main blades postgenetics development microform frost heaving and frost sorting.

Figure 48. The solifluction terrace formed on the left side of upper part of Khushuut valley, complicated by the micro-relief of the tufurs (left) and its cut (right). Sutai ridge, a.s.l. 2675 m.

Figure 49. Open forms of cryogenic landsliding in the upper part of Tsagangol valley, Sutai ridge, a.s.l. 2780 m.

The internal structure covered solifluctional forms revealed the cycles of their development. It is established that each cycle of the soil flow was completed by the stage of soil formation. The thickness of the soil horizon of the present stage indicates a longer period of its formation, compared with the two older.

Open forms of cryogenic landsliding are formed due to the primary processes of freezing-out of larger debris on the surface during of the day from sediments with a dominant loamy fraction, followed by a downward displacement. Open forms have a low scarp (20-40 cm) and limiting its border, in scheme of having the form of tongue.

When moving up the slopes, their natural rejuvenation is noted, expressed in the recency of morphological elements and the absence of lichen coating on the debris. Modern formations numerous, but much smaller in size than the older. The specificity of the structure of permafrost forms indicates the existence in the recent past of conditions conducive to the more intensive development of cryogenic processes.

A distinctive feature of the periglacial zone of the Sutai ridge is the absence of lakes, which is explained by the topographical features of the territory, and the morphology of glaciation caused by them: the predominance of glaciers of the hanging type, which do not produce frontal-moraine complexes, creating favorable conditions for the formation of periglacial lakes.

Figure 49.  Spatial distribution of glacial lakes in the Tsambagarav massive. 1 – The position of glaciers in 2017, 2 – the area of deglaciation from Little Ice Age, 3 – glacial lakes

In the belt of recent deglaciation of the Tsambagarav ridge, was noted the appearance and increase in the area of the waters of 8 glacial lakes, in contrast to the Sutay massif. The vast majority of lakes which formed within fifty years dedicated to foreground modern glaciers (moraine-dammed lake and lake of intermoraine depressions) and glacial-accumulative complexes of the Little Ice Age (thermokarst lakes). The periods of increased activity of glacial and thermokarst limnogenesis (1992-98, 2012-2016) were revealed by means of geoinformation analysis of polychronous spatial data.

 

Study and collect of pollen of higher plants and spores of the lower ones in Altai Mountains

 

The study of modern pollen the additional basis for reconstructing past climate and vegetation [Zhang et al., 2017]. At the present time many studies are focused on relationships between pollen and climate, especially on local and regional scale. Therefore, in this project we study modern pollen spectra obtained in Altai Mountain. In 2017, the samples were collected in Tauber traps, established in accordance with the requirements of the Pollen Monitoring Program on the territory of the Altai State Biosphere Reserve. These samples were analyzed by the Scanning Electron Microscope Hitachi S-3400N (Japan) in IWEP SB RAS. Now in IWEP SB RAS are the collection more than 120 species of the most common plants that grow in different part Altai Mountains. In this project the pollen of each plant was photographed from various angles for a more accurate spatial representation of the object, namely in the polar and equatorial positions (figure 50).

Figure 50. Photos of Cordalic nobilis (a, b) and Populus tremula (c, d) pollen from various angles

Figure 51. An approximate view of the Atlas pages

In this project, it is planned to create an Atlas of pollen of Altai plants — woody, shrubby and grassy forms. In the Atlas there will be photos of pollen under different magnifications and different foreshortening, photographs of vegetative forms of plants or herbarium material, maps of distribution areas of species, a botanical description of the species and a morphological description of pollen grains. An approximate view of the Atlas pages is shown in Figure 51.

Conclusion

 

The Planetary climate experienced significant changes during the 20th century. Global changes consist in raising the main characteristic of the Earth’s climate — global temperature. Modern climatic changes are clearly demonstrated in all regions of the Earth. The average air temperature has increased by 0.74 °C, since the beginning of the 20th century, and about two-thirds of this growth has occurred since the 1980s. Each of the last three decades was warmer than the previous one. The air temperature was higher than in any previous decade, since 1850. [America’s Climate Choices. — Washington, D. C.: The National Academies Press, 2011. — P. 15].            Warming was accompanied by climatic anomalies everywhere, as a result of which regional climates underwent significant changes, which was most clearly expressed by the beginning of the XXI century.

The analysis of climatic changes (temperature and precipitation) in Altai Mountains showed step change points, namely a significant temperature increase in 1996 and precipitation in 1980. The most significant increase in temperature when comparing two intervals (1957-1995 and 1996-2017) occurred in autumn (1.28⁰C) and spring (1.13⁰C), while for annual and summer the average temperature increase was approximately the same (0.86). The maximum increase in mean precipitation from the first (1957-1979) to the second (1980-2015) period was obtained for the annual and summer (more than 25 mm), a smaller increase for spring (5 mm). The analysis of medium-period observations showed that the average annual air temperature in Western Mongolia increased by 2.07 °C.

Climatic changes have led to significant and irreversible changes in the spatial structure of the nival-glacial and cryogenic systems of the highlands of the Mongolian Altai. The area of deglaciation total increased by 37.5 sq. km, from the time of the maximum of the Litlle Ice Age (about 18 sq.km in the last 50 years) on the mountain ranges of Sutai and Tsambagarav. Landscape belts moved due to climatogenic uplifting to a height of 180-200 m in the mountain-glacial basins of the Mongolian Altai. Subglacial deposits became subaerial, under the influence of a changed climatic background, in the deglaciation zone of the ridges. The newest periglacial and limno-periglacial landsystems began to be formed on a young lithogenic basis.