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Physical Geography of Northern Eurasia

Climatic Change and the Development of Landscapes

<<< Cenozoic Climates: The General Trend | Physical Geography Index | Climate at Present and in the Historical Past >>>

Synthesis: The Development of a Zonal Structure in Modern Landscapes

The evolution of climates and landscapes of Northern Eurasia was marked by an extremely wide range of transformations in the Cenozoic. The range of climatic and landscape changes recorded during the last 60 Ma was the largest in the landscapesphere history of the planet. These transformations were mainly controlled by global changes in heat balance although the state of the Earth's surface also influenced the global climate system.

The character of climatic fluctuations changed through time. At the higher thermal level of the early Cenozoic, the climate system experienced long-period fluctuations (5-10 Ma) with an amplitude of 2-4∞C (Velichko, 1989). The end of this interval was marked by the Eocene optimum, the principal thermal maximum of the Cenozoic. Between the end of the Eocene and the Pliocene, the frequency of the oscillations increased, their period reduced to 1-5 Ma, and the amplitude of mean annual temperatures increased (Figure 2.15).

Mean January temperatures in the Cenozoic in four regions of Northern Eurasia

Fig. 2.15 Mean January temperatures in the Cenozoic in four regions of Northern Eurasia. Compiled by A. Velichko and V. Nechaev

About 1 Ma BP, climatic changes became more regular and quasi-cyclic in character. The alternating glacial and interglacial phases formed macrocycles. The duration of each macrocycle varied between 200 Ka to 100 Ka, decreasing progressively. Temperature also decreased from the earlier macrocycles towards the later ones. Thus temperature during an optimum of each of the three best expressed interglacials, which occurred in the last half a million years, was about 1∞C lower than the temperature of the previous optimum. The same tendency for an increasing severity of climate has been observed during the glacial phases. For example, while only ground veins developed in the central East European plain during the cold phases of the early Pleistocene, ice wedges formed in the second half of the middle Pleistocene. This is a weighty evidence of permafrost existence (Nechaev, 1996). The maximum expansion and the lowest temperature of permafrost occurred during the cold phases of the late Pleistocene.

Temperature changes in time, therefore, show a certain regularity: slow and rather smooth 'long-wave' oscillations were typical of the early Cenozoic; 'middle-wave' rhythmical fluctuations with growing amplitude occurred in the Neogene; and high frequency cyclic fluctuations developed in the Pleistocene. As thermal levels were in general declining, fluctuations became more pronounced and more clearly recorded in sediments (Sinitsyn, 1980b). The colder the climate became, the stronger the effect its fluctuations (even the small ones) had on landscapes, especially in higher latitudes, where changes in mean annual temperature by 2-4∞C induced considerable shifts in zonal boundaries. High frequency oscillations at a lower thermal level were amplified by positive feedbacks such as the ice-albedo feedback (Velichko, 1973). The decrease in the mean global temperature was accompanied by an increase in the equator-pole temperature difference, which changed from 10-12∞C at the beginning of the Cenozoic to 40-42∞C in the Holocene.

Within the framework of the Meso-Cenozoic megacycle, the modern stage is a cold stage of the late Cenozoic and a warm (interglacial) phase of the latest Holocene macrocycle. The interglacial began at 10.3 Ka BP and its optimum occurred at 6-5.5 Ka BP. The present time, therefore, belongs to the second half of the interglacial and is characterized by climatic cooling.

Climatic change controlled the development of the zonal structure (Figure 2.16). For a long time, Northern Eurasia was a single vast forest zone (or hyperzone). The internal structure of this zone changed in time.

Schematic representation of the dynamics of latitudinal zonality during the Mesozoic and Cenozoic

Fig. 2.16 Schematic representation of the dynamics of latitudinal zonality during the Mesozoic and Cenozoic. After Velichko (1989)

Throughout the Paleocene and the first half of the Eocene, subtropical evergreen forests covered western and southern Eurasia and had a high proportion of tropical elements in the south of Central Asia (paratropical sub-zone). The northern part of the continent was occupied by the mild temperate forests with abundant subtropical plants. The impoverished subtropical forests survived in the south through the late Eocene and Oligocene. The mild temperate forests still remained in the mid-latitudes and in Kazakhstan and Central Asia. In the forests of Western Siberia, small- and broad-leaved species coexisted, while in higher latitudes mixed broad-leaved coniferous forests developed for the first time. Later, temperate mixed forests expanded southwards. By the middle of the Miocene, savannah-like landscapes had developed in Kazakhstan, and later in the south of the East European plain and Western Siberia. A completely new zonal component Ч steppes Ч developed at the end of the Miocene. In the Pliocene, the tundra biome formed in response to further cooling. Semi-deserts and deserts developed as a result of increasing aridity.

These conclusions are in good agreement with a previously developed concept of planetary zonal pattern evolution (Velichko, 1989) which suggested that in the Paleogene and Neogene a monozonal (or hyperzonal) structure was replaced by a polyzonal one. Until the late Pliocene, changes in zonal structure were 'migratory' (i.e., changes referred to shifts of zonal boundaries).

The dynamics of biomes in the Quaternary were different. A new type of hyperzone, the cryohyperzone, dominated by open landscapes (such as periglacial steppes, tundra and dry steppes) developed in response to the severe climates of glacial epochs, while the forest zone became much smaller. During the warm phases, a polyzonal structure, in particular the forest zone, was restored. Such a regime of alternating polyzonal and cryohyperzonal structures may be described as a 'transformational' or 'pulsating' type of zonal dynamics.

Although a polyzonal structure was restored during every interglacial, zonal plant communities were different from those of the previous warm phases. Similarly, landscape zones were not restored precisely at the same positions. For example, tundra almost disappeared during the earlier interglacials, only surviving as a narrow strip along the Arctic coast in the Holocene. The type of the modern polyzonal system is that of a cooling of the post-optimal stage of an interglacial. This system includes elements of very different ages. Thus tundra communities developed in the Eopleistocene whereas steppe communities date back to the Miocene. In other words, the modern zonal structure is heterochronous.

Longitudinal differentiation of biocoenoses developed with time. It was very poorly expressed in the early Cenozoic when subtropical forests covered both the East European plain and Siberia. Since the late Oligocene, and especially in the middle Miocene, landscapes became more differentiated. Thus landscapes of north-eastern Asia became well adapted to cooling and later to increasing aridity. Climatic change in this region could be forced by neotectonic uplifts: the newly forming mountains prevented the transportation of relatively warm and moist air by the westerlies (Sinitsyn, 1965). However, the formation of a topographic barrier was not the only factor. Under conditions of progressive planetary cooling, a high pressure system (the Siberian anticyclone) developed over the mountains and became one of the most important atmospheric centres. With mountain building, its cooling influence spread further westwards and contributed to the formation of sea ice in the Arctic. Another high montane system, the Himalayas-Tibet, formed at the same time with another anticyclone of global importance being related to it (Li, 1991).

A radical alteration took place in the atmospheric circulation and climatic system of Northern Eurasia in the Cenozoic. Throughout most of the Paleogene, westerlies were of primary importance under the conditions of high heat supply and subdued topography. In the Neogene (in particular, since the middle Miocene) and Quaternary, active orogenesis, which occurred in the south and east of the region, significantly changed the atmospheric circulation. It may be concluded, therefore, that the modern climatic system began to form about 15 Ma BP and became similar to the modern one ,in the Pleistocene.

Modern landscapes, the landscapesphere as a whole, and individual landscape units combine components of different age. First, there are considerable differences in the ages of the biotic and abiotic components. Plant communities are the youngest and respond to changes most quickly. They are different from those of the Holocene optimum. At present, they adapt to general cooling. Soils are more conservative and some soil profiles integrate processes of the whole Holocene history. Abiotic components are older as a rule and may bear imprints of ancient events. Thus, in the formerly glaciated regions, strong differences between the types of the surface exist, the morainic relief being the youngest type within the limits of the latest Pleistocene ice sheet expansion. Rather young are the subsurface deposits in the loess regions where loess accumulation terminated just before the Holocene. However, the late Pleistocene permafrost impacted surface topography: the relict cryogenic microrelief remained after the ice wedges melted in late glacial time. The microrelief, in turn, affects the development of soils and erosion at present.

To conclude, a typical feature of modern landscapes and landscapesphere is heterogeneity and many of their properties and components have been inherited from the past.

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