The present distribution of permafrost on the Qinghai-Xizang (Tibet) Plateau (QTP) is largely a relict of the permafrost formed during the late Pleistocene. It has been degrading and shrinking in areal extent under the fluctuating climates, with a general trend of warming, during the Holocene. The major criteria for the occurrence of relict permafrost include the remnants of ancient buried permafrost, relict permafrost tables, thawed sandwiches (taliks), thick-layered ground ice, and periglacial phenomena such as pingo scars, cryoturbations, primary sand and clayey silt wedges, ice wedge casts, aeolian sand dunes and loesses, thick layers of peat, and humic soils. On the basis of 14C dating of soils, comprehensive analyses, and comparisons of the spatiotemporal distribution of relict and modern permafrost and periglacial phenomena, the evolution of permafrost and periglacial environments since the late Pleistocene was divided into seven stages: (1) the cold period at the end of the late Pleistocene (35,000 to 10,800 years B.P.); (2) the period of significant climatic change during the early Holocene (10,800 to ∼8500–7000 years B.P.), (3) the Megathermal period in the middle Holocene (∼8500–7000 to ∼4000–3000 years B.P.), (4) the cold period in the late Holocene (∼4000–3000 to 1000 years B.P.), (5) the warm period in the later Holocene (1000 to 500 years B.P.), (6) the Little Ice Age (500 to 100 years B.P.), and (7) the recent warming period (100 years B.P. to present). The conditions for permafrost development, distribution, and the paleoclimates and paleoenvironments are discussed for each stage.
If you can't find a tool you're looking for, please click the link at the top of the page to "Go to old article view". Alternatively, view our Knowledge Base articles for additional help. Your feedback is important to us, so please let us know if you have comments or ideas for improvement.
 The Qinghai-Tibetan Plateau (QTP), bounded by the Pamier (Pamir) Plateau in the west, by the Bayankala (Bayan Har) Shan and Hengduan Shan Mountains in the east and from the Kunlun Shan, Aerjin Shan (Altun, or Altyn Tagh), and Qilian Shan Mountains in the north to the Himalayas in the south, is an area of numerous east-west trending mountain ranges and contiguous basins (Figure 1). It has an approximate area of 2.58 million km2 and is the source of three major rivers: the Huang He (Yellow), the Chang Jiang (Yangtze) and the Lancang Jiang (Mekong) rivers in southeastern Asia. The QTP is seasonally buffeted both by the monsoons from the southeast and south, and by the dry winds from central Asia. Vegetation varies from subtropical forests in the southeastern part of the QTP to alpine (high-cold, or paramos) deserts in the western area, but generally much of the Plateau is that of a high plains pastoral landscape.
 On the basis of the presence of brownish and yellowish relict-related crusts and pollen records, there were more humid and warmer climatic episodes which almost completely thawed previous permafrost layers prior to the Last Glaciation. Most Chinese scholars now agree that the main body of permafrost on the QTP was formed during the Last Glaciation of the late Pleistocene [Zhou, 1965a; Wang et al., 1979; Ding and Guo, 1982; Wang et al., 2000; Zhou et al., 2000; Jin et al., 2006] after most of the QTP essentially had reached its present general elevation. They also agree that during the following Holocene (about 10,800 years B.P. to present), several periods of permafrost development and partial decay occurred in response to fluctuating changes in climate, although the climates overall remained periglacial. Zhou [1965b] believed that the deeper (15–20 m) residual permafrost from the Last Glaciation in the interior QTP was connected vertically with the most recent permafrost from the Neoglaciation (3000–5000 years B.P.), but the belief in that connection has been challenged [Wang, 1989; Li and Wang, 1992]. The developments of and the degradations of the permafrost along the periphery regions of the QTP have been much more sensitive to small changes in temperature when compared to those from similar changes in temperatures occurring in the interior.
2. Study Region
 This paper discusses the evolutionary history of permafrost in the central and eastern Qinghai-Tibetan Plateau (QTP) since the end of late Pleistocene, using relict permafrost and periglacial phenomena along the Qinghai-Tibet Highway (QTH) from Golmud to Lhasa, the Qinghai-Kang (western Sichuan) Highway (QKH) from Xi'ning to Yushu (Figure 1), adjacent areas, and the Xinjiang-Tibet Highway (XTH) from Yecheng to Lhasa. The relationships between latitudinal and elevational permafrost during the Last Glaciation Maximum (LGM), and the boundaries of relict permafrost, the implications thereof, are briefly discussed.
 The general characteristics of the QTP geomorphology include (1) high elevations (greater than 4500 m in the interior and 3000 m on the peripheries); (2) interior mountain ranges, high plateaus, intermontane basins alternating with wide valleys in northwestern to southeastern or westerly to easterly directions; (3) higher in the northwest and lower in the southeast; and (4) surrounded by high mountains with deeply incised valleys. The major geomorphological units include from north to south: Aerjin Shan Mountains (Altun or Altyn Tagh) to Qilian Shan Mountains, Qaidam Basin, southern Qinghai to northern Tibetan Plateau, southern Tibetan Plateau and the Himalayas. Numerous high mountains and deep gorges occur on the eastern and southeastern edges and in the Hengduan Shan Mountains.
 The rapid uplifting of the QTP during the Quaternary greatly influenced the environments on the Plateau, its contiguous regions, and eastern, southern and central Asia. Although there have been numerous studies, and the timing of the QTP uplift still is poorly known and lively debated, there appears to be general agreement that the Plateau had reached its present elevation and general configuration by the late Pleistocene [Li et al., 1979; Li and Fang, 1999]. Recent oxygen-isotope-based estimation of the paleoaltimetry of late Eocene and younger deposits of the Lunpola basin in the center of the Plateau indicate that the surface of Tibet has been at an elevation of more than 4000 m for the past 35 ± 5 Ma [Rowley and Currie, 2006]. The authors have accepted this estimation, and the general agreement among the various authors that the Plateau had reached its present elevation by the late Pleistocene, in the preparation of this paper. The uplifting of the QTP casts profound impacts on the formation and evolution of the permafrost. The high elevations and large areal expanse provided the crucial low temperatures for the development of the permafrost.
 The 1150-km-long Qinghai-Tibet Highway (QTH) traverses representative permafrost areas in the interior (Figure 2). The 650-km section from Xidatan (Highway Maintenance Squad Station, or HMSS 60) to Liangdao'he (HMSS 125) is in the permafrost area. Mostly continuous permafrost occurs along the 560-km-long section from Kunlun Shan Mountain Pass to the Amdo'nan'shan Mountains (near HMSS 116). From Xidatan to Kunlun Shan Mountain Pass, and from the Amdo'bei'shan Mountains to the Liangdao He River, island permafrost occurs at elevations above 4125–4250 m and 4640–4680 m, respectively. The complex and active tectonic activities also have resulted in major geomorphology differentiating permafrost development and evolution. As a result, permafrost is relatively thin, and warm with numerous taliks. In the interior, alpine permafrost is thicker and colder than that on the plains and high plateaus. However, the thermal stability is complicated because of the combinations of moisture, surface canopies and resultant differences in soil thermal conductivity.
 The primary focus of this paper is the evolution since the late Pleistocene of permafrost along the QTH in the interior (southern Qinghai Plateau and northern Tibetan Plateau) where presently subarctic semiarid climatic conditions prevail at elevations generally of 4500–4800 m on the plateaus and 5000–6000 m in mountainous areas. The pertinent data from along the QTH, QKH and Xinjiang-Tibet Highway (XTH), and from scientific expeditions to other areas of the Plateau, indicate that the mean annual air temperatures (MAATs) range from −0.6 to −9.8°C and the annual precipitation from 20 to 500 mm.
 The western QTP is dominated by alpine (high-cold or paramos) desert landscapes, the eastern area by grasslands and meadows. More than 70–80% of the interior QTP has continuous permafrost where the MAATs generally are colder than −4°C (Table 1) and the observed mean annual ground temperatures (MAGTs) range from −0.1 to −3.5°C (Figure 2). Sporadic, or island permafrost as warm as +0.5 to +1°C in the MAGTs, and usually indicating degrading permafrost and/or changing surface or subsurface conditions, surround the areas of continuous permafrost.
Table 1. Division and Major Characteristics of Permafrost in the Eastern QTP
Distribution of Permafrost
Areal Continuity, %
Alpine island permafrost on the northeastern QTP
Middle and high mountains, alternating with valleys and basins, elevation ∼3500–4000 m, relative height ∼500–1000 m, with flat and wide intermontane basins
Permafrost (quasi-stable, transitory, unstable, very unstable) Seasonally frozen ground
Continuous permafrost in the Burhanbuda Shan and Animaqin Shan Mountains
Higher than 4500 m, steep topography in the east, with relative heights of about 1000 m; in the sources of the Yellow River, hills prevail with relative heights of 200–500 m, and flat and wide topography.
Continuous, thickness and MAGTs greatly affected by locality.
Permafrost (transitory, unstable, very unstable) Seasonally frozen ground
 The observed thicknesses of permafrost range from 10 to 175 m, but can be highly variable, especially in areas of island permafrost generally less than 20 m in thicknesses. The observed depths of measurable annual change in ground temperature, the zero amplitude of change, vary from 10 to 16 m. There are three situations in the continuous permafrost zone. In hilly or alpine areas where the MAGTs range from −1 to −4°C, the permafrost thicknesses range from 30 to 130 m; in the high plateau river valleys of the sources of the Chang Jiang (Yangtze) River, where the MAGTs vary from −1 to −1.5°C, permafrost generally is thinner than 60 m; and in other river valleys with MAGTs from 0 to −1°C, permafrost is generally less than 50 m. The MAGTs in talik zones beneath large rivers generally range from +0.5 to +1.0°C and are dependent on both width and to some extent depths of the rivers.
 Along the QTH, the lower limit of permafrost (LLP) is at 4150–4250 m in the north and 4450–4560 in the south. In the western Kunlun Shan Mountains, continuous permafrost is found from 10 km south of Dahongliutan (elevation 4450 m) to 30 km north of Domar (4630 m), a total distance of 331 km along the XTH; island permafrost is then encountered [Li et al., 1998]. In the eastern QTP, alpine permafrost occurs in the Burhanbuda (Burhan Budai) Shan, Animaqin (A'nyêmaqên) Shan and Bayankala (Bayan Har) Shan Mountains of the eastern Kunlun Shan Mountains, with LLPs at 3840–3900 m in the Heka'nan'shan Mountains, 4000–4050 m in Animaqin Shan Mountains, 4150–4200 m in Bayankala Shan Mountains, and at 4250–4300 m in Yushu County, Qinghai Province. For every 1° southward, the rise in the LLP is 120–130 m along the QTH (32°–36°N, 92°–94°E), 130–145 m along the QKH (33°–37°N, 97°–100°E), but only about 80 m along the XTH (34°–37°N, 78°–80°E) [Zhou, 1965b; Zhou et al., 2000]. The LLP also decreases eastward. For example, it is at 4500 m at the Kalakunlun (Karakorum) Shan and western Kunlun Shan Mountains, 4300–4350 m in the Muzitag (Muztag) Shan Mountains (87°–88°E, 36°–37°N) [Li et al., 1998], 4200 m at Xidatan, and 3900 m in the Heka'nan'shan Mountains; that is, for every increase of 1°E, the LLP declines by about 27 m. In the Kunlun Shan Mountains, the MAGTs decrease with elevation at rates of 7 to 8°C/km (elevation) from Xidatan to Kunlun Shan Mountain Pass, while the average values on the QTP are 5–8°C/km. The thickness of permafrost increases with elevation at rates of 200 m (thickness)/km (elevation) along the QTH and 100–200 m/km along the QKH.
3. Study Methods
 The determination of relict permafrost boundaries using relic periglacial phenomena have been widely applied in Europe, the former USSR, and in North America. For example, the relict southern limit of permafrost (SLP) in Europe during the Würm glacial stage of the late Pleistocene was determined using sand wedges and the accompanying cryoturbations/involutions. Soviet geocryologists established the evolution of permafrost in the former USSR since the late Pleistocene using the evidence from polygonal wedge structures, pingo scars and thermokarst lowlands. The relict SLPs at each stage were estimated based on the fact that the distribution of polygonal ice wedges was 5° to 6°N of the present southern limit of permafrost (SLP) in western Siberia and 2° to 4°N in the central Asian mountains [Baulin, 1981]. Among the relict periglacial phenomena, polygonal wedge structures are the most reliable evidence of permafrost occurrence, but their indicated environments vary greatly. Other phenomena such as thermokarst depressions, pingo scars, rock fields and rock glaciers, involutions and cryogenic terraces also can suggest the previous presence of permafrost, but their reliability is compromised by their wide range of developmental environments, subject to interpretations, and the difficulties in identifying their origins. However, if cross examined with other paleoclimatic indicators, they can provide convincing results.
 Most of the dating for permafrost relied on the 14C ages of the peat, humus, or other materials rich in organic carbon, found in either epigenetic or syngenetic permafrost. Because the soils had been formed during (syngenetic permafrost) or before (epigenetic permafrost) freezing, the age of humic soils indicates the greatest possible age (syngenetic permafrost) of the frozen soils. If they are epigenetically frozen, then the age of permafrost would be younger than the age of the humic soils. Humic soils themselves are not the indicators of permafrost. Rather, they might indicate the thawing of permafrost during the warm periods. However, they generally contain more organic carbon, which is generally used for the carbon dating. Frozen humic soils can indicate permafrost, because they are in the continuous permafrost zone. If they are outside the permafrost areas, and without other periglacial indicators such as involutions, ice wedge casts, and others, it might be difficult to infer whether it is an evidence of relict permafrost, or it was formed in a nonpermafrost region. The accelerator mass spectrometry (AMS) 14C dating was used by some cited authors. The ages dated using thermoluminescence (TL), optically stimulated luminescence (OSL) and infrared-stimulated luminescence (ISL) were used by others. Relative dating of soil strata is very important for dating relict permafrost elements and some periglacial remains, and thus is carefully studied for interpreting the evolution of permafrost.
 The climatic records from ice cores, tree rings, plant and animal remains, and lacustrine sediments can be compared and contrasted with reference to the occurrence of permafrost or periglacial environments, and can provide persuasive deductions regarding the boundaries of relict permafrost. It is generally assumed that the internal forces that shaped or have been shaping the permafrost environments on the Qinghai-Tibetan Plateau (QTP) have been relatively stable during the discussed periods, and that the responsive behaviors of permafrost to paleoclimatic changes are similar to those of today. Also, it is quite difficult to reconstruct relict permafrost processes due to thermal inertia, slow response to climatic changes, and marked regional variability. Calculations indicate that significant changes of permafrost temperatures at depth might require hundreds to thousands of years to respond to surface heat and mass exchanges. Only the coldest or warmest stages would leave recognizable “footprints” in the permafrost; the “memories” of secondary climatic fluctuations would be erased during the following more extreme conditions. The existing permafrost generally does not exceed 150 m on the QTP, and the main body of permafrost apparently was formed during the late Pleistocene. In this paper, seven distinct stages of permafrost evolution since the late Pleistocene LGM are discussed.
3.1. Deeply Buried Permafrost and Relict Permafrost Table
 Thick-layered, repeatedly segregated ground ice is generally found at the permafrost table on the QTP, and many works have been published on its features [Cheng and Wang, 1983] or the formation mechanisms [Cheng, 1983]. It was generally formed in fine-grained soils (silts, clays, fine sands, or gravels with fine grains well mixed). It is usually a soil-ice mixture, with volumetric ice content more than 50% and soil particles floating in the ice, forming the ataxitic or breccia-like cryostructure.
 With more data acquired, and analyses using new research methods, it is now evident that permafrost on the QTP responded to climatic warming during the Megathermal in the Holocene and the Neoglacial cooling during 3000–5000 years B.P. In the late 1970s, analyses of soil cores from several tens of boreholes along the Qinghai-Tibet Highway (QTH) indicate that in the permafrost regions, segregated ice layers concentratively occur at three depths: 2–4 m (the modern permafrost table), 7–8 m and 14–16 m [Cheng and Wang, 1983]. It is generally agreed that regional occurrence of multilayer segregated ice concentrated at certain depths can indirectly indicate the permafrost table [Kudryavtsev, 1978]. Therefore the widespread detection of segregated ground ice layers at depths of 7–8 and 14–16 m in permafrost on the QTP may indicate two relict permafrost tables, respectively.
 Since the 1980s, one or two layers of deeply buried permafrost have been encountered at elevations from ∼4200 to 4400 m in boreholes or water wells in warm (≥−1°C) permafrost areas in the eastern QTP (Figure 3), during hydrological and engineering surveys and exploratory drilling. The deeper layer is buried at depths of 15–20 m, with residual permafrost ∼10–20 m in thickness. In some boreholes, or wells, a 4- to 5-m-thick upper layer of permafrost also is encountered at a depth of about 10 m. There is a 5- to 7-m-thick layer of thawed soil (talik) between these two layers. For example, in Borehole ZK8, the layer from 0 to 1.8 m in depth is seasonally frozen, the layer from ∼1.8 to 11.6 m is thawed, that from ∼11.6 to 15.2 m is the upper layer of permafrost, that from ∼15.2 to 20.0 m again is a talik, and that from ∼20.0 to 31.2 m is the lower layer of permafrost. Boreholes CK1 and ZK2, and the Qingshui'he water well have similar cryolithology with Borehole ZK8. However, only the deeper layer of residual permafrost was encountered in Borehole ZK2, the southernmost borehole in the permafrost area along the QKH. In Boreholes CK5 and ZK6, the permafrost is connected to the active layer, suggesting that significant decaying of the upper permafrost layer has not been occurring at those locations during recent years. In the Huashixia water well, however, the upper layer of permafrost from a depth of 5.3 to 8.2 m is disconnected from the active layer at a depth of 2.7 m by a 2.6-m-thick talik. Between the upper and lower (at ∼16.3–19.3 m) permafrost layers, there is a talik resulting from an ongoing decay of the upper layer. The lower layer of permafrost developed in an area with lacustrine or wetlands soils, such as ice-rich silty clays and clayey silts, and in Borehole ZK6, there is an ice layer between 19.8 and 24.3 m. The combinations of seasonal frost, and upper and lower talik layers and permafrost reveal important information on climatic changes during different periods.
 When the permafrost table was maintained for a long time, such as in an equilibrium state, a contrasting distribution of soluble salts and unstable minerals would occur above and below the permafrost table. In a typical rock weathering profile, the weathering degree and content of soluble salts and unstable minerals decrease with depth. When the weathering soils are frozen, the mobility of and accumulation of soluble salts resulting from chemical weathering, are greatly reduced below the permafrost table. Therefore, compared to that prior to the freezing of the soils, they have not changed significantly. However, above the permafrost table, the eluviation is still active, and weathered soluble salts are removed continuously. Unstable minerals contents are thus reduced. Using this hypothesis, the positions of relict permafrost tables were reconstructed based on the contents of soluble salts obtained from the boreholes samples, the β value curves for clay minerals, and the δ value curves for clastic minerals. Additional positions were based on samples from boreholes at sections from the Kunlun Shan to Fenghuo Shan Mountains. (The β value, an indicator for unstable mineral contents, is the ratio of peak strength of hydromica to that of chlorite. In the Quaternary sediments on the QTP, hydromica and chlorites were the major clay mineral combinations. Hydromica is stable, but chlorites are unstable in weathering processes. The β value can therefore indicate the depth of weathering. The peak strengths of mineral were obtained using a diffractometer, with an error less than 5%. The δ value, an indicator of weathering, is defined as the ratio of stable and extremely stable minerals contents to relatively stable and unstable mineral contents of heavy minerals finer than 0.25 mm.) It is concluded that there are two positions of relict permafrost tables, at depths of ∼7–8 and ∼14–16 m [Xing and Ou, 1983]. Xing and Ou  studied relict permafrost tables through the analyses of the variations of soluble salts (salinity) and clay mineral contents, and found that the contents of soluble salts spiked at the abovementioned three depths. This phenomenon was interpreted as the result of accumulation of soluble salt from more active soil layers above or at the permafrost table.
 Using mathematical-physical methods, Ding and Guo  simulated the history of permafrost evolution during the Holocene. They assumed that during the Neoglaciation the MAATs ranged from −4.6 to −7.8°C, while the modern MAATs varied from +0.8 to −2.4°C from Xidatan to the Tanggula Shan Mountain Pass. The results from mathematical-physical simulations indicate that if the MAATs were 0.8 to −2.4°C from Xidatan to Tanggula Shan Mountain Pass during the middle Holocene Megathermal (∼8.5–7.0 ka B.P.), a 15-m-thick layer of permafrost could have thawed along the QTH. If they were −4.6 to −7.8°C during the late Holocene colder periods, a 30-m-thick layer of soils could have refrozen during the colder periods of the late Holocene [Ding and Guo, 1982]. That is to say, the thawed permafrost would have refrozen and connected with the previous permafrost layers. These results agree well with the relict permafrost tables found at the depths of ∼14–16 m and the lower layer of residual permafrost which had remained in the eastern QTP during the Megathermal [Xing and Ou, 1983]. During ∼1000–700 years B.P., there was a warm period with MAATs about 2°C warmer than at present when permafrost degraded downward [Xie and Zeng, 1983]. During this warmer period, the relict permafrost table retreated to the depth of ∼7–8 m along the QTH; the upper layer of residual permafrost remained on the eastern Plateau during the downward degradation.
3.2. Relict Permafrost and Ground Ice
 Relict permafrost and ground ice provide the most direct evidences of permafrost evolution and climatic changes since the late Pleistocene. Recent surveys revealed double or single layer deeply buried permafrost in the sources of the Huang He (Yellow) River and on the northern slopes of the Bayankala (Bayan Har) Shan Mountains on the northeastern QTP, such as in Borehole ZK2 at Qingshui'he and Borehole ZK6 on the northern bank of Eling (Nygoring) Hu Lake in Maduo (Madoi) [Wang, 1989]. Several sites along the QTH also have been identified for dating the ages of permafrost formation:
 1. Borehole 8 at Xidatan is near the northern lower limit of permafrost (LLP). The present thickness of permafrost is about 20 m. Six layers of humic soils are visible at depths of 1.85, 2.50, 3.75, 4.40, 7.10, and 8.58 m. In the lowest two layers, undecomposed plant roots and stems are larger than existing counterparts, indicating a more humid and warmer climate favoring plant growth and the accumulation of organic matter. The dating of humic soil at a depth of about 4.40 m indicates that it was formed about 7530 years B.P. [Cui, 1980].
 2. Permafrost profile north of HMSS 109 is at a knickpoint west of the QTH on the northern slope of the Tanggula Shan Mountains. The age of the frozen humic soils at a depth of 2 m in the permafrost is 5058 ± 443 years B.P., indicating the maximum age of the permafrost at this location.
 3. A permafrost island 1.5 km north of HMSS 120 has 8- to 10-m-thick permafrost. The ages of humic soils at depths of 0.8 and 1.9 m are 4363 ± 178 years B.P. and 4576 ± 648 years B.P., indicating that the permafrost was formed after 4350 years B.P.
 4. Permafrost islands at the two sides of a gulch 1 km north of HMSS 121 near Liangdao'he have a total permafrost area of about 0.4 km2. According to pit excavations, the permafrost table is located at the depth of ∼1.0–1.1 m, and the bottom of the permafrost is at 1.7 m, i.e., the permafrost is only ∼0.6–0.7 m in thickness. Undecomposed plant remains, with a 14C age of 780 ± 131 years B.P., are contained in the humic soils at the depth of 1.6–1.7 m, i.e., permafrost was formed during the past several centuries, perhaps near the beginning of the Little Ice Age (LIA).
 The relict ice of segregation origin in permafrost on the QTP is generally found in the vicinity of the permafrost table; and its widespread presence required long-term ample water supply and moisture migration [Cheng, 1982]. It is closely related to climatic fluctuations and the resultant changes in the positions of the permafrost tables. Therefore concentrative distribution of ground ice, especially the thick-layered ice in the Quaternary lacustrine sediments, can indirectly indicate the position of the relict permafrost table.
 1. Borehole CK113-2 is in a small intermontane basin southwest of HMSS 113 (near Amdo) along the QTH. In the strata, clayey soils, sandy soils, fine sands and silts occur alternately, with the permafrost table at about 2.9 m in depth. There are three sections of relatively thick ice layers, indicating three periods of ground ice formation. During the first period, alternating ice layers (∼3–5 cm in thickness) and silty clay were formed at ∼15.6–17.8 m. During the second period, ∼2- to 5-cm-thick pure ice layers were formed at ∼7.8–8.9 m. During the third period, layered or coated ice was formed at ∼3.0–4.8 m.
 2. The permafrost table is at 1.9 m in Borehole 203 at Qingshui'he along the QTH. The lower section of the thick-layered ice is from 16.1 to 19.5 m in depth, with the thickness of most layers varying from ∼5 to 8 cm with the largest 12 cm. The middle section is at ∼7.8–8.1 m, the maximum thickness of the ice layers is about 9 cm. The upper section is at ∼1.9–4.2 m, and the ice layers are about 1–2 cm in thickness.
 The common features of ice formation at the abovementioned locations include: (1) the ice layer thickens and purifies downward in the profile, suggesting an origin of repeated accumulation, ample water supply, and favorable water migration conditions; (2) three periods of ice formation are apparent, with similar distribution depths at the various boreholes. The upper layer is related to recent permafrost formation. The middle and lower layers of ground ice coincide with positions of the two relict permafrost tables, corresponding to the formation periods of the two layers of relict permafrost [Xing and Ou, 1983].
3.3. Pingo Scars and Depressions Formed by Collapsed Pingos
 There is a series of depressions (about 50) formed by collapsed pingos at elevations from ∼4250 to 4300 m on a front part of a late Pleistocene glacio-fluvial fan on the water divide between Dongdatan and Xidatan. This front part of the fan is located on a recent fault (7530 ± 300 years B.P.) [Cui, 1980]. These pingo scars are protruding, volcano-crater-like, and horseshoe-shaped in the planar view. They generally occur on the edges of the fault basin, where faults controlled the edges of the basin. The Xidatan Valley, which is a fault basin, has been subsiding intensively since the late Pleistocene. The relationship between intensive subsidence and pingo development is only indirect, but the subsidence is favorable for the formation of syngenetic permafrost, and development and evolution of relict permafrost and ground ice are closely related with the geological background. The continuous and intensive subsidence might have helped in the formation of ground ice layers near the permafrost table.
 There are landform evidences of pingo collapsing, such as volcano-crater-like microtopography, horseshoe-shaped formations, water accumulation in the center and around the pingo scars; and ground water outflows and subsequent peat accumulation. The diameter of the depressions is about 100 m. The relative depths at the centers are about 5–6 m. At present, most of the depressions have vegetative coverage. The age of silty clay from the center of one of the depressions is 720 ± 39 years B.P. The preliminary conclusion from the well-preserved remains of the collapsed pingo, and analyses of the MAATs required for developing such large-scale pingos, is that they were probably formed during the cold period of the Little Ice Age (LIA). The lower limit of permafrost (LLP) at present at this location is ∼4400–4450 m, which is about 150–200 m higher than that during the LIA period; the MAATs were about 1°C colder than at present assuming a 6.5°C km−1 lapse rate of atmospheric temperature.
 There are groups of pingo scars on the northern side of K65 along the Maqin to Changma'he Highway. The depressions formed by collapsed pingos are distributed in a string of connected ponds at elevations of about 3800 m on the lower parts of piedmont slopes. There are other accompanying features related to the depressions identified as pingo scars, such as volcano-crater-like microtopography, horseshoe-shaped formations, water accumulation in the center and ground water outflows. The age of humic soils at the center of one of the depressions is 625 ± 57 years B.P.; that at a depth of 20 cm on the top of the periphery ridge is 3925 ± 185 years B.P. Therefore the existence period of the pingo group is ∼3925–625 years B.P. The present LLP is above 4100 m; the mean annual air temperatures (MAATs) were at least about 2°C colder than at present.
 There also is a group of pingo scars at about 4, 000 m in elevation, 40 km east of Shiqu (Sêrxü) County Town, Sichuan Province. The 14C age of organics in the fine sands and silts at 10 cm in depth on the periphery ridges is 3295 ± 175 years B.P., which suggests that they formed almost simultaneously with the abovementioned two groups of pingo remains. This location is now in the seasonally frozen ground zone. Because the LLP nearby is about 4200–4300 m, the MAATs were about 1.5 to 2°C colder than at present assuming a 6.5°C km−1 lapse rate of atmospheric temperature with elevation.
3.4. Cryoturbation and Involution Remnants
 Cryoturbations, or involutions, from frost heave and thaw settlement in fine-grained soils on gentle slopes produce folded and disturbed cryogenic structures under the influence of gravity [Harris, 1981]. Since ice segregation and soil creep are involved in the development of cryoturbation, its remnants are reliable evidence of permafrost occurrence [Van Vliet-Lanoe, 1985; Coutard and Mucher, 1985].
 Cryoturbation remnants in many areas of present seasonally frozen ground zones have been discovered on the QTP. The major characteristics include intensively folded structures, and inversed structures, at shallow depths with thicknesses from 1.5 to 2.0 m. They generally have conspicuous inconformity with overlying and underlying strata. The disturbed soil profiles indicate earlier relict permafrost.
 Involutions 39,830 ± 3840 years B.P. in age were discovered from the late Pleistocene proluvium in the sources of the Huang He (Yellow) River [Cheng et al., 2006]. They also are found together with other cryogenic structures, particularly wedge structures.
 In a peat valley at Qi'nongga, north of Yangbajing (Yangbajain) (30°04′26″N, 90°33′10″E, 4080 m) along the QTH, the thickness of peat is about 4 m. The 14C ages of organic soils 0.3 and 1.6 m below the top of the peat layer are 3270 ± 70 and 6130 ± 90 years B.P., respectively. The remnants of cryoturbation structures are found in alternating gray clayey silts and yellow sands and gravels above the peat layer. Therefore the cryoturbation developed later than 3270 years B.P. [Li, 1982]. The 14C age of the ash-like sandy soils containing peat in the strata on the first terrace of the Kunlun He River southeast of Nachitai (Naij Tal) is 4910 ± 100 years B.P. [Pu et al., 1982]. The cryoturbation remnants visible (Figure 4) at the abovementioned two places belong to the ∼4000–3000 years B.P. cold period. On the basis of the difference of present and relict LLPs, the MAATs were about 2.0–2.5°C colder then than today.
 The age of the bottom of the thick peat layer is 9970 ± 135 years B.P. in the profile at the 4200-m elevation along the Wuma Qü River near Dangxiong (or Damxung, 30°29′N, 91°06′E, 4200 m). Layered cryoturbation structures formed during the early Holocene (10,800 years B.P.) are visible below this peat layer. The present LLP is ∼600–700 m higher than that during the early Holocene, indicating that MAATs then were about 4°C colder assuming the lapse rate of atmospheric temperature at 6.5°C km−1 (elevation).
 The Ruo'ergai (Zoîgé) Peatland in the eastern QTP developed at elevations from ∼3410 to 3500 m where only seasonally frozen ground is now found. On the basis of hand pit excavations, the maximum frost penetration is about 1.2 m in the wetlands on the second terrace along the southern bank of the Hei'he River. The age of peat at 1.0 m is 22,600 ± 750 years B.P. The disturbance and folding are intensive in soil profiles at the knickpoints. This reveals that permafrost was very well developed on the Ruo'ergai (Zoîgé) Peat Plateau by the end of late Pleistocene. Permafrost is now found only in mountainous areas higher than 4500–4600 m, or 1000 m higher than the relict LLP. Therefore the MAATs at the end of the late Pleistocene were about 6°C colder than today.
3.5. Primary Sand Wedges and Ice Wedge Casts
 Primary sand wedges and ice wedge casts are reliable indicators of the occurrence of relict permafrost. However, environmental conditions for their formation, and thus the indicated soil or air temperatures, vary greatly. In fine-grained soils, ground ice can form and volumetric heaving result at temperatures several degrees below 0°C. However, thermal contraction occurs in fine-grained soils only in environments colder than −5 to −6°C [Mackay, 1974]. Ice wedges are formed by the repeated contraction and cracking of soils and accompanying freezing of soil moisture and snowmelt year after year [Lachenbruch, 1962]. When frost cracks are filled by sands, primary sand wedges are formed. An ice wedge cast results from sand and soils filling the void resulting from the melting of an ice wedge [Black, 1976]. On the basis of field observations and laboratory experiments on frost cracking, and development from ice to sand wedges, it is evident that they are closely related to existing soil types, moisture contents, and ground temperatures, particularly the cooling rates of winter temperatures [Lachenbruch, 1962]. Sand wedges can be formed at about −2.0 to −1.0°C in ground temperatures in clays, silts and peat, and −5.0 to −4.0°C in medium or coarse sands or gravels. However, ice wedges also can be formed at −8 to −4.0°C, or colder [Romanovskii, 1977].
 Many wedge structures have been discovered on the QTP during the past few decades [Xu and Pan, 1990; Pan and Chen, 1997]. A number of sand and soil wedges have been identified at various locations, generally located in the alluvial strata on the second terrace of majors rivers, and overlain by deposits ∼1–2 m in thickness. Collapsed structures, and fallen gravels, apparently related to the secondary filling by existing sediments after melting of the ice, are visible in some ice wedge casts.
3.5.1. Wedges in the Northeastern QTP
 The northeastern QTP refers to the area from 34° to 38°N and from 110° to 114°E, with elevations from ∼3000 to 4400 m, which includes Qinghai Lake, Gonghe and Xinghai basins, and the sources of the Huang He (Yellow) River. About 80 wedges of various types have been identified [Xu and Pan, 1990]. The shapes, filling materials, existing soil types and deformations, and types of wedge structures are included in Table 2.
Table 2. Characteristics of Sand Wedges in the Northern QTPa
 On the Ruo'ergai (Zoîgé) Plateau, at 3400 to 3500 m in elevation, fossil cryoturbations and ice wedge casts filled with silts provide evidence of relict permafrost [Lehmkuhl, 1995; Lehmkuhl and Hövermann, 1996]. A 1.4-m-thick layer of sandy silt covers the wedges. The basal layers of this silt were thermoluminescence (TL) dated to be 18 ka B.P. [Lehmkuhl, 1995]. Permafrost exists with a wide range of MAATs from −3 to +1°C in the nearby mountains. The present MAAT is about 0.7°C, and the lower limit of permafrost (LLP) is at 4300 m. Assuming the lapse rate of atmospheric temperature at 6.5°C km−1 (elevation), the cooling was more than 6°C during the formation period of the ice wedges because the lowering of the LLP was close to 1000 m.
 A late Pleistocene sand wedge, 0.7 m in width at the top and 1.2 m in height, was investigated at Qiejitan (36°17′N, 101°09′E), at the 3100-m elevation in the Gonghe Basin. The wedge was filled with medium to coarse sands and clayey soils with near-vertical bedding, and surrounded by similarly aged soils without visible deformations. Two layers of paleosols with loess in between were found on top of the wedge structure. Because of the lack of deformation in the surrounding soils, vertical bedding inside the wedge, and the relatively small size of the wedge, it is believed that it is a primary sand wedge. The age of the upper paleosol is 19,430 ± 360 years B.P.; that of fill materials at the bottom of the wedge is 20,403 ± 430 years B.P. [Pan and Chen, 1997]; and the sand wedge predates the paleosols. The MAGTs during the sand wedge formation are believed to be about −3.0°C [Romanovskii, 1977]. The differences between the MAATs and MAGTs presently range from 2.5 to 3.0°C. Therefore the MAATs during the wedge formation period were −5.5 to −6.0°C. The MAATs at the location presently are about 1.0°C, and it can be inferred that the MAATs during sand wedge formation were about 6.5–7.0°C colder than at present.
 A group of ice wedge casts is located on a gentle slope in Xingxiuhai (34°50′N; 98°00′E; ∼4200–4300 m), south of Maduo (Madoi) in a source of the Huang He (Yellow) River [Xu and Pan, 1990]. About 50 wedge structures were identified in a roadcut. The wedges are ∼0.5–1.6 m wide at the top and ∼0.8–2.0 m in height. The surrounding rocks are deeply weathered, intensively deformed shale. Densely cemented medium to coarse sands fill the wedges; occasionally, some boulders or gravels also were included. The 14C ages of the CaCO3 contained in the medium and coarse sands vary from 12,300 ± 100 and 13,490 ± 1430 to 16,340 ± 245 years B.P. [Pan and Chen, 1997; Cheng et al., 2006], while a TL date of about 18 ka B.P. also was reported [Lehmkuhl and Haselein, 2000; Lehmkuhl et al., 2000]. The MAGTs when the ice wedges were formed would have been about −7.0 to about −6.5°C. Compared with the present MAGTs of about −0.5 to 0°C, the MAATs during the ice wedge formation should have been about 6.0 to 6.5°C colder than at present.
 Ice wedge casts younger in age also were identified in sandy gravel beds on the second terrace of the Huang He (Yellow) River in the same region [Cheng et al., 2006]. They are characterized by higher ratio of breadth to depth and are 0.8–1.4 m in breadth and 0.5–0.9 m in depth. On the basis of the TL dating, they were formed at the middle Holocene (5690 ± 430 and 5430 ± 410 years B.P.). Cheng et al.  described two major types of ice wedge casts, and indicated that “they occur in groups” in the exposed sections on the northern bank of Eling (Ngoring) Lake, and at Yematan, Huanghe Town, Mashaling, Maduo (Madoi) County, and north of the Ye'niu'gou Valley. One type of wedge structure, found in sandy gravel beds on the second terrace of the Huang He (Yellow) River, was characterized by a higher ratio of breadth to depth, and is 0.8–1.4 m in breadth and 0.5–0.9 m in depth. The lower border of the ice wedge cast is rounded in section. The other type, discovered in bedrocks on the second terrace, has a lower ratio of breadth to depth, and is 0.3–1.0 m in breadth and 0.7–2.0 m in depth. Its lower border is sharp. Cheng et al.  concluded that they are all ice wedge casts. They did not say how many ice wedge casts were identified, but there were only two involutions identified. On the basis of the TL dating, the former was formed during the mid-Holocene (5690 ± 430 and 5430 ± 410 years B.P.), and the latter during the deglacial period (about 12–15 ka B.P.), which is correlated to the oldest Dryas in age. The involutions were discovered from the late Pleistocene proluvium, and are 39,830 ± 3840 years B.P. in age. Since the tensile strength of rock is greater, and it breaks more cleanly, than “icecrete” composed of ice and sandy gravels, the data given suggest that the climate was much colder during 39,810 years B.P. than during 5600 years B.P.
 Ice wedge casts in bajada gravel beneath loess of late and postglacial age, with TL dated at 15,100 ± 1600 years B.P. (LGM), also were discovered near the northern base of the Qinghai'nan'shan (southern Qinghai Shan, or Xiangpi Shan) Mountains (36°34.85′N, 100°29.67′E, 3305 m) [Porter et al., 2001]. In several nearby exposures, the upper 2 m of bedded bajada gravel was involuted. However, this was the only ice wedge cast reported there.
 Thirty-nine sand wedges, with top widths of 12–40 cm and with depths of 17–80 cm, were discovered in at late Pleistocene alluvial gravel sites in many areas in the Hexi Corridor (37°–40°N) north of the Qilian Shan Mountains. The filling material was fine aeolian sand. These sand wedges are overlain by Holocene gravel or loess. The AMS 14C dating indicates that their formation ages vary from 19,100 ± 125 to 22,500 ± 190 years B.P. (LGM, Marine Isotope Stage (MIS) 2), when temperatures were 11–13°C colder than at present [Wang et al., 2003].
3.5.2. Wedge Structures Along the Qinghai-Tibet Highway (QTH)
 The soil and sand wedges found along the QTH are representative of those in the interior Qinghai-Tibetan Plateau (QTP). The soil wedges at Nachitai (Naij Tal) (35°53′N, 94°30′E, ∼3550–3600 m) are now in the seasonally frozen ground zone. The wedges are in the bluish and grayish clays and the yellow sandy clays on the second terraces of the major tributaries of the Kunlun He River. The soil wedges are ∼0.3–0.4 m wide at the top, 1.1–1.2 m in depth and are surrounded by cryoturbation remnants with vertical beddings. Loess-like clayey soils “inharmoniously” overlay the cryoturbation structures and wedges. The fill is similar to the overlying strata. The ages inside the soil wedges and of the surrounding soils are 14,041 ± 399 and 15,377 ± 292 years B.P., respectively, indicating that these soil wedges were formed near the end of late Pleistocene. The difference in ages suggests that the ice wedges were formed and had continued to exist for about 1300 years before a warming climate melted the ice allowing the more recent overburden to slump into the voids created. On the basis of the analyses of the surrounding soils, the MAGTs for the development of the soil wedges would have been about −2°C and the MAATs about −5.0 to −4.5°C colder than at present. A more recent loess layer at a depth of 0.4 m was dated at 7207 ± 387 years B.P., suggesting an average loess accumulation of about 0.6 cm per century since 14 ka B.P.
 There is a sand wedge group in the sand and gravel on the second terrace along the eastern bank of the Zuomaoxikong Qü River on the northern piedmont of the Fenghuo Shan Mountains. The elevation is 4660 m and the location 34°46′N, 92°55′E. The top widths of the sand wedges are 1.5 to 1.8 m, the heights are 1.5 to 1.7 m, and are overlain by a layer of sands and gravels 30 to 50 cm in thickness. The 14C age of the grayish fine sands containing humus at the bottom of the No. 1 sand wedge is 23,500 ± 1200 years B.P. The grayish silts in the middle section of the No. 2 sand wedge have a 14C age of 15,340 ± 770 years B.P. (Figure 5). The brownish fine sands overlying the wedges is 9218 ± 189 years B.P. [Zhang, 1983]. The data suggest that there was a warmer climate about 23,000 years B.P. allowing for accumulation of organics, followed by a colder period, and then a warmer climate about 15 ka B.P. again allowing for the accumulation of organics. The melting of the ice wedges, and the subsequent slumping of the overburden into the voids created, would appear to have occurred sometime between 15 and 9.2 ka B.P.
 Sand wedges generally do not exhibit very much indication of deformation in the horizontal bedding of the surrounding sediments, while ice wedges and ice wedge casts show a very definite upward movement of the adjacent soils as the surrounding soils attempted to reexpand into the void created by the tension crack. The tension cracking may be a periodic (even decades) and rare phenomena, depending on suddenness of winter, it severity, and sufficiency of moisture in the sediments to provide their cohesion. Both sand and ice wedges tend to occur in groups. However, the sand wedges are more common in drier sediments where the snow cover may be thinner during some winters and the winds can fill the voids created by the tension cracks with wind-blown silts, fine sands and vegetative debris. Sand wedges tend to be more widely spaced, and their bottoms more oval-shaped than the bottoms of ice wedge casts. Sand wedges also may exhibit a saw-tooth shape on the side due to the prevailing wind direction when that vertical segment of the filling material was deposited, or some vertical bedding if material is washed in.
 The sand wedges at HMSS 82 (34°37′N, 92°48′E, 4710 m) are in the sand and gravel strata on the second terrace of the Chi Qü River along the southern piedmont of the Fenghuo Shan Mountains. Eight sand wedges have been identified in one group along the 100-m-long knickpoint of the second terrace. There is a layer of alluvial and fluvial gravely sandy clay 0.3 to 0.8 m thick. The age of gray sands containing humus inside a wedge is 9160 ± 170 years B.P. This suggests that these sand wedges were formed during the same period (from the end of late Pleistocene to early Holocene) as those mentioned above. Similar sand wedge groups have been discovered on the second terraces along the Tongtian He, Tuotuo He and Bu Qü rivers in the interior QTP. According to the soil analyses, the sand wedges along the southern and northern piedmonts of the Fenghuo Shan Mountains were formed under environments with MAGTs about −6 to −5°C. The present MAGT is about −1.0°C, i.e., assuming unchanged differences between MAGTs and MAATs, when the sand wedges were formed the MAATs and MAGTs were 4 to 5°C lower than today.
3.5.3. Ice Wedge at Tianshuihai, Western QTP
 Tianshuihai (35°15′N, 79°50′E, 4850 m) is on the southern slope of the western Kunlun Shan Mountains. The inactive ice wedge (Figure 6), along the knickpoint of the Akesaiqin (Aksayqin) Lake shore, is the only existing inactive ice wedge so far identified on the QTP. The top width is 0.7 m, the depth is 1.5 m, and the ice veins contain significant amount of clayey soils [Li and He, 1990; Li and Jiao, 1990]. The surrounding soils are grayish clay. The contained clayey soils were aged 15–21 ka B.P.; the ice formation could be no earlier than these ages.
 The difference between the MAGTs and MAATs can be assumed to be about 2.5 to 3.0°C on the QTP [Li and He, 1990; Li and Jiao, 1990]. Romanovskii  concluded that colder MAGTs (<−6 to −7°C) were required for ice wedge formation in silty soils. Péwé  believed that active ice wedge formation MAATs range from −6 to −12°C, inactive ice wedges from −2 to −6°C, and that the growth of ice wedges needs −6 to −8°C. Washburn  stated that ice wedge growth needs air temperatures of −5 to −8°C. Some geocryologists believe that the frost cracking is more closely related to the MAGTs: Washburn  noted that in eastern Siberia, frost cracking occurred in frozen ground at −4 to −5°C; Mackay  observed that in the Mackenzie Delta, ice wedges were cracking in frozen ground at −1 to −2°C. Some other people believe that frost cracking occurs when the low temperatures at the upper permafrost layer (5–10 m), or at the permafrost table, are lowered to colder than −15 to −20°C, and are very dependent on the rapid cooling rates. If it is quick enough, frost cracking can occur in frozen ground at −4°C [French, 1976]. North American scholars believe that precipitation is not really a factor in the formation of ice or sand wedges because, in Alaska and northern Tibet where precipitation is less than 100–200 mm, ice wedges can occur; in Siberia, sand wedges have been identified in areas with precipitation of 300–400 mm. Therefore the ice wedges were formed at a paleotemperature of about −8.5 to −9°C as the soils were similar to those reported by Romanovskii . While scholars from North America have reported frost cracks forming at widely differing temperatures [Péwé, 1963; French, 1976; Washburn, 1979; Mackay, 1986], most have concentrated on MAGTs apparently without much reference to soil types, or to the rapidity of the freezing [Lachenbruch, 1962].
3.5.4. Sand Wedges in the Keliya (Keria) River Delta, Northwestern QTP
 Sand wedge groups associated with 2- to 3-m-diameter polygonal networks on wide alluvial fans were identified along the margin of the sandy desert in the delta of the Keliya (Keria) River (between 36°–39°N and 81°–83°E) in the northwestern margin of the QTP, at elevations of 1760 and 1660 m [Yang et al., 2006]. These wedges, filled by aeolian sands with distinct vertical laminae, are 0.5–1.0 m in depth and up to 40 cm in width. OSL ages of the aeolian sands filling the wedges are 40,600 ± 3200 years B.P. (MIS 3) at 1760 m and 18,300 ± 1300 years B.P. (LGM, MIS 2) at 1660 m, probably indicating depressions of MAATs by 11–18°C [Hövermann and Hövermann, 1991; Yang et al., 2002, 2006].
3.6. Aeolian Sands, Loesses, Thick-Layer Humus, and Peat
 The aeolian sands and loesses on the western and central QTP apparently were formed under arid climates in environments with sparse vegetation. Thick-layered humic and peat soils on the eastern QTP were formed under warmer and more humid monsoon climates with relatively denser vegetation.
3.6.1. Aeolian Sands
 According to the strata and surface manifestations, the major deposition of aeolian sands on the Plateau since the late Pleistocene can be divided into three stages:
 1. Stage 1 is the end of the late Pleistocene: The surface aeolian sand landforms from this period have been losing their original shapes. The 30- to 80-cm-thick sand layers are widespread in basins, on valley bottoms, and on the high plateaus. The sand layers can be up to 3 m in thickness in knickpoint profile as revealed in a river bank north of Wudaoliang. The age of the humic soils is 12,700 ± 820 years B.P. at the depth of 1.0 to 1.1 m.
 2. Stage 2 is the early Holocene: The surface manifestations of aeolian sands include alternating distribution of fixed and semifixed sand dunes, sand bars and aeolian lowlands. The remains of plants and humus in the sand strata conspicuously increase upward in the profile as in a wind-eroded sand ridge along the lake shore 2 km east of Wudaoliang. Fine sands containing plant remains are present in alternating layers about 2.5 m in thickness. The age of the plant remains at this site is 9716 ± 270 years B.P. at the depth of 1.8 m. The upward increase in organic content suggests a warming climate with increasing precipitation during this period.
 3. Stage 3 is the recent period: The aeolian erosion on the Plateau has been quite intensive recently, with well-developed mobile sand dunes and ridges generally without vegetation. Recent aeolian sand deposits and landscapes are widely observed on the southern slope of Xidatan'bei'shan Hill, along the southern bank of the Hong He River, on the eastern bank of Hong Hai Lake in Anduo (Amdo) County, and in the Gonghe Basin on the northeastern QTP. Land desertification burying highways, farmlands and grasslands has been accelerating under the arid and warming climate since the LIA.
 Loesses and aeolian sands were deposited concurrently. By the end of the late Pleistocene, the extent of continuous loess deposition reached the northern piedmont of the Bayankala (Bayan Har) Shan Mountains. Loesses also were well developed and have been preserved at the higher ends of the third terraces and in the fluvial fans on the northern slopes of the Kunlun Shan Mountains. The thickness of loess, with well-developed vertical joint structure, reaches more than 10 m on the third terrace of the Kunlun He River near the Ge'ermu (Golmud) Reservoir. The age of the loess is 18,931 ± 400 years B.P. at the depth of 10 m. The age of loess inside the soil wedges in Nachitai (Naij Tal) is 15,377 ± 292 years B.P. The age of loess deposits on the second terrace of the Tongtian He River south of Qumalai (Qumarlêb) is 12,475 ± 182 years B.P. These 14C dating data suggest that there was a warmer time, about 15 ka B.P., during the relatively long period of cold and arid climates near the end of the late Pleistocene.
3.6.3. Thick Layers of Humus and Peat
 The age dating of the peat and humus layers on the QTP indicates that the majority of the organic soils were formed during the middle Holocene Megathermal, beginning about 8.5 to 7 ka B.P. and ending about 4 to 3 ka B.P. The age data on a north-south transect along the QTH are as follows:
 1. The age of the sandy clays, containing ash, on the first terrace at Nachitai at 4910 ± 100 years B.P. might be evidence of ancient human activities.
 2. In the Borehole 8 soil profile at Xidatan (35°43′N, 94°13′E, 4460 m), there are six layers of humic soils at depths from 3.75 to 8.58 m, in which the age of the humic soils at the depth of 4.4 m from the surface is 7530 ± 300 years B.P.
 3. In the vertical permafrost profile on the southern slope of the Tanggula Shan Mountains north of HMSS 109, the humus layers are as thick as 3.5 m. The age of the humic soils is 5058 ± 443 years B.P. at the depth of 2 m from the surface.
 4. The humus is 2.5 m thick in a permafrost island at HMSS 120 north of Anduo (Amdo). The ages of humic soils at depths of 0.8 and 1.9 m from the surface are 4363 ± 178 and 4576 ± 648 years B.P., respectively, suggesting a continuous warmer period during the deposition.
 5. The peat layer is as thick as 4 m in the Qinongga valley north of Yangbajing (Yangbajain). The ages of the peat are 3270 ± 70 and 6130 ± 90 years B.P. at depths of 0.3 and 1.6 m from the surface, respectively.
 The strata data (Figure 3) from the northeastern and eastern areas of the QTP have two consistent characteristics: they indicate that (1) the middle Holocene Megathermal had thawed the late Pleistocene permafrost to depths of 15 m in the easterly and to 20 m in the more easterly areas and (2) that the climate within those areas during the Megathermal was warmer with longer growing and significantly wetter summer seasons resulting in the development of the thick layers of peat. Age dating of some of the peat soils includes the following:
 1. The middle section of 2-m-thick humus on the eastern slope of the Riyue Shan Mountains is 4920 ± 80 years B.P.
 2. The peat layer on the southern slope of the Heka'nan'shan Mountains is about 5 m in thickness. The age at a depth of 2 m is 4625 ± 117 years B.P.
 3. The age of the humus layer in the lobe of a solifluction fan in the Wenbo'nan'shan Mountains, Shiqu (Sêrxü) County, Sichuan Province is 4395 ± 215 years B.P.
 4. The peat layer on the Gadangsong Glacial Moraine (33°20′N, 101°08′E, 4154 m) in the Nianbao'yuze Shan Mountains in Jiuzhi County, Qinghai Province is 4.15 m in thickness. The age of the bottom peat layer is 5422 ± 94 years B.P.
 5. During 9350 to 370 years B.P., a peat layer 5.2 m in thickness was deposited at the Peat Farm on the outskirts of Hongyuan Town on the Ruo'ergai (Zoîgé) Plateau, in which a 3.3-m-thick peat layer suggests continuous deposition from 6350 to 3250 years B.P. [Sun, 1998].
 Large quantities of data indicate that during the middle Holocene, the climate was relatively warm and wet, favoring plant growth and humus accumulation, and the formation of thick layers of humic and peat soils. The paucity of aeolian sands and loesses within the peat layer also suggests the relative absence of barren lands from which to obtain aeolian materials during this period.
 The main body of existing permafrost on the QTP was formed during the Last Glaciation in the late Pleistocene [Zhou et al., 2000]. Therefore the reconstruction of the evolution of permafrost has to begin from the late Pleistocene. On the basis of the comparisons and contrasts of the distribution of recent and relict permafrost, and periglacial phenomena, the evolution of permafrost since the end of the late Pleistocene can be divided into seven stages (Table 3): (1) the cold period (35,000 to 10,800 years B.P.) at the end of the late Pleistocene, (2) the period of significant climatic changes during the early Holocene (10,800 to 8500 years B.P.), (3) the Megathermal period (8500 to 4000 years B.P.), (4) the late Holocene cold period (4000 to 1000 years B.P.), (5) the later Holocene warm period (1000 to 500 years B.P.), (6) the Little Ice Age (500 to 100 years B.P.) and (7) the recent warming period of the past century. The largest or smallest areal extent of permafrost during each stage is estimated by assuming similar relationships among the permafrost distributions, MAGTs, and MAATs during those periods with those of today.
Table 3. Evolution of Permafrost and Cold Regions Environments on the QTP During the Holocene
Start Time, ka B.P.
Relict Permafrost and Periglacial Phenomena
Paleoclimates and Paleoenvironments
Comparison and Contrast With Present LLP, and Estimated Total Areas of Permafrost
End of the late Pleistocene
Last Glaciation Maximum
Sand and soil wedges, ice wedge casts, cryoturbation, loesses
Cold and arid climate, MAATs ∼5.0–6.5°C colder than today, glaciers, permafrost and periglacial phenomena well developed
The LLP 1000 m lower than today, thickness of permafrost of about 200 m, continuous and extensive permafrost, with a total area 70–80% greater than today.
Significant climatic changes
Sand and soil wedges, aeolian sand dunes and ridges
Unstable climate, but with a warming and humidifying trend, ∼3.0–4.0°C colder than today in MAATs, permafrost and periglacial.
The LLP ∼600–700 m lower than today, stable and widespread permafrost, with total area 40–50% greater than today.
Megathermal, or Hypsithermal
Relict permafrost table, lower layer permafrost, thick layer ground ice, thermokarst lakes and lowlands, thick layer peat and humus
Warmer and more humid climate with MAATs ∼2.0–3.0°C warmer than today, and development of wetlands reached its climax
Intensive permafrost degradation, thawing downward to depths of ∼15–25 m, LLP ∼300–400 m higher than today, continuous permafrost survived only in mountainous areas, other areas only in islands or deeply buried, with total permafrost area ∼50–60% less than today.
Neoglaciation cold period
Pingo groups, polygons and stone circles, cryoturbation, moraines and tills
Relatively cold and dry periglacial climate, 2.0°C colder than today, permafrost and periglacial phenomena well developed
Epigenetic permafrost formed by refreezing thawed soils, connected vertically in the interior, taliks existed in marginal areas, LLP 300 m lower than today, with total area ∼20–30% greater than today
Climatic warming, MAATs ∼1.5–2.0°C higher than today, thermokarst lakes well developed
Relative degradation of permafrost, downward thawing reached ∼10–15 m in depths, LLP ∼200–300 m higher than today, with total permafrost area ∼20–30% less than today
Relict stone fields, streams, mobile sand dunes and bars, permafrost islands, moraines and tills
Relatively cold and dry climate, MAATs ∼1.0–1.5°C colder than today, many small and middle size salt and saline lakes
Refreezing of thawed soils to a depth of about 10 m, distribution of permafrost similar to present, the LLP ∼150–200 m lower than today, with total permafrost area about ∼10–15% greater than today
Upper layer talik and permafrost, warming of permafrost, increasing thickness of the active layer
MAATs rose by ∼0.3–0.5°C during the past 40 years, conspicuous retreat of permafrost and glaciers, and shrinkage of river runoffs, lakes, and wetlands, degeneration of grasslands, and expansion of deserts
Extensive permafrost degradation, LLP rose by ∼40–80 m, total area of permafrost declined by ∼6–8%
4.1. Cold Period at the End of the Late Pleistocene
 During the glaciations from 330 to 220 ka B.P., permafrost was continuously distributed [Shi et al., 1992a]. Pan and Chen  identified at least four periods of permafrost expansion during the past 150 ka. The Last Glaciation Maximum (LGM) occurred during 30 to 25 ka B.P., climates were shifting from those of warm, semiarid steppes to cold and arid periglacial environments [Hövermann, 1998]. On the basis of the analyses of sand, soil, and ice wedges (casts) in Tianshuihai, Nachitai (Naij Tal), the Fenghuo Shan Mountains, Gonghe Basin, Hexi Corridor, and in the sources of the Huang He (Yellow) River [Wang, 1989; Wang et al., 2003; Cheng et al., 2006]; cryoturbations or involutions on the second terraces of major rivers such as Kunlun He, Tuotuo He, Tongtian He and Bu Qü rivers [Guo, 1979; Qiu, 1982; Cheng et al., 2006]; cryogenic terraces in the Keliya (Keria) area [X. P. Yang et al., 2002, 2006, 2004]; and palynological temperatures [Guo, 1976; Tang and Wang., 1976]; the MAATs would have been 5 to 7°C, even 10°C, colder than today, i.e., −8 to −11°C. The MAGTs would have been colder than −4 to −5°C in the western and central Plateau, with continuous permafrost 120 to 150 m in thickness. The lower limit of permafrost (LLP) extended down to 1800 m, and the snowline to 4000 m. In the eastern QTP, however, the continuity of the permafrost was disrupted by deeply incised gorges and gradually transited to alpine permafrost. In the northeast, the northern LLP at 1950–2300 m was connected with that in the Qilian Shan Mountains and was probably connected with the latitudinal southern limit of permafrost (SLP) [Zhou et al., 2000; Cui et al., 2004].
 By the end of the LGM, climates were gradually warming. However, permafrost still was widespread except in the deep valleys in eastern and southeastern Tibet. In the north, the LLP was near the Ge'ermu (Golmud) Reservoir (∼3100–3200 m in elevation). In the south, it was close to Yangbajing (Yangbajain) (∼3900–4000 m). In the east, the expanse of permafrost extended into the Gonghe Basin (∼2400–2500 m in elevation) and Shiqu County (∼3400–3500 m), Sichuan Province. Permafrost also was widespread on the Ruo'ergai (Zoîgé) Plateau. The LLPs were about 1000 m lower than today, and the total area of permafrost, including glaciated areas, was about 70 to 80% greater than today. Permafrost was extensively developed, except for the taliks formed due to the presence of abnormally high geothermal heat flows, and beneath large rivers and deep lakes. Permafrost was thick, cold and thermally stable.
 There is evidence of the presence of permafrost in the Chaidamu (Qaidam) Basin during the LGM with cold, windy and dry climates, with subsequently enhanced desertification and extensive aeolian deposits [Xu et al., 1982; Xu and Xu, 1983]. The high-level lakes, encompassing Cha'erhan (Charhan, or Qarhan) and Chaka (Caka) lakes, formed during the interglacial stadial (50 to 32 ka B.P.), disaggregated into smaller lakes. The sharp reduction in lake areas, and salinization occurred during 24 to 18 ka B.P. Many polygonal sand and silt wedges were formed along lake shores and beneath drained lakebeds. Sand wedges, aged about 18,500 years B.P., developed on the higher lake terraces (about 3000 m in elevation) [Chen, 1981]. The Chaidamu Basin was in the cryolithozone during the LGM. However, because of the high salt contents, permafrost was largely in the form of cryopegs. Aeolian sands and loesses were extensively deposited because of enhanced erosion and transportation under the cold and arid periglacial environments. Many areas are still covered by a layer of fine sands more than 0.5 m in thickness. The 3-m-thick sand layer north of Wudaoliang and the sand dunes on the second terrace in the Gonghe Basin were formed during this period. Ice, silt and sand wedges, cryoturbations, and polygonal structures also were well developed. The loesses on the northern slopes of the Kunlun Shan Mountains are believed to have been deposited during this period. The water supplies for lakes had declined sharply, and evaporation had increased, resulting in higher salinity and hardness of lake waters. The large quantity of evaporates in the Cha'erhan Salt Lake area in the Chaidamu Basin started to form about 24,000 years B.P. Carbonate enrichment also is evidenced by the presence of large quantities of carbonate nodules in lacustrine sediments in the Qingshui'he area along the QTH. The 14C age of silicified plant roots in the lacustrine sediments of a lake on the northern bank of the Tuotuo He River is 14,810 ± 350 years B.P.
 The climate was the driest during ∼16–9 ka B.P. and sylvites were deposited in dried salt lakebeds [Chen, 1981]. On the northern Plateau, permafrost extended to the periphery of the Talimu (Tarim) Basin with the LLP at 1900–2900 m and the SLP at 39°N (Figure 7). During the post-LGM period, some ice wedge casts developed in the sources of the Huang He (Yellow) River and along the northern bank of Qinghai Lake.
4.2. Abruptly Changing Climate During the Early Holocene (10,800 to 8500–7000 years B.P.)
 With the elevation of the QTP above 4000 m by the late Pleistocene, large-scale glaciations gradually ceased. The climate in the early Holocene was very unstable, with abrupt changes. For example, the δ18O records from ice cores from the Dunde Ice Cap in the Qilian Shan Mountains indicate that there was a minimum value of −12.75‰ at 8700 years B.P., which might suggest a sharply cooling period, and a maximum value of −9.6‰ during ∼8500–8400 years B.P., which marks a possible warming event [Yao, 1992]. In general, the trend of climatic change during the early Holocene was a fluctuating increase in temperatures and wetness.
 According to the distribution of relict periglacial phenomena, the large expanses of permafrost existing at the end of the late Pleistocene started retreating during this period. The lower limit of permafrost (LLP) on the margins of the QTP rose by ∼300–400 m. The northern LLP rose to ∼3400–3500 m at Nachitai (Naij Tal), while the southern LLP elevated to ∼4200–4300 m between Yangbajing (Yangbajain) and Dangxiong (Damxung). Generally, the LLPs were ∼600–700 m lower and the MAATs were ∼3–4°C colder than at present. Although permafrost was degrading, it was thick, extensive and continuous, and its total area still was ∼40–50% greater than today.
Tang  analyzed the distribution of vegetation based on pollen records recovered from lacustrine sediments during the early Holocene. The vegetation was dominated by mesophytic deciduous broadleaf and coniferous forests in the south during ∼10,000–9100 years B.P.; and it was subalpine meadow dominated by Artemisia with a cool and dry climate in the Qinghai Lake area during ∼11,000–8000 years B.P. It was a grassland environment with a cool and moist climate in the west during ∼10,000–7700 years B.P. [Tang, 2000]. Under a general warming and humid climate, wetlands were formed in valleys and basins, depositing peat and thick-layer humus. Thick layers of peat with bottom ages of 8175 ± 200 and 9970 ± 135 years B.P. started to form in Wumaqü and Qinongga [Li, 1982]. The blackish silty clays at depths of ∼2.5–3.0 m (soil samples from a depth of 2.95 m) were dated at 8800 ± 305 years B.P. in Borehole CK80-3 in Qingshui'he along the QTH. It then graded from yellowish clayey soils, containing limestone blocks and carbonate nodules, to fine and medium sands, and to blackish silty clayey soils in the lower section of the borehole profile, indicating a warming and wetter climate as evidenced by the increased accumulation of organic matter.
 The age of the silts and sands containing humus in the upper part of sand wedges on the second terrace of the Zuomaoxikong Qü River on the northern slopes of the Fenghuo Shan Mountains is 9218 ± 189 years B.P. [Zhang, 1983]; and the age of humus-containing silts and sands is 9160 ± 170 years B.P. inside the sand wedges at HMSS 82 on the southern slope. These suggest an abating of frost action, and the ceasing of frost crack growth. The frost cracks filled with sands, gravels and soils, and sand wedges stopped developing, indicating a warming and wetter climate favoring plant growth, and the developing of fixed and semifixed sand dunes at the end of the late Pleistocene. The presence of the multilayer undecomposed plant roots (aged 9716 ± 270 years B.P.) in a sand ridge 2 km southeast of Wudaoliang also provides supporting evidence.
 There are uncertainties about the wetness during ∼9500–8700 and ∼7200–6300 years B.P., when the lakes on the QTP were several times larger than today according to the relict shorelines. More than one third of the northern QTP could have been covered by lakes [Avouac et al., 1996], and it can be inferred that deep thawed basins existed beneath those lakes [Brewer, 1958], particularly beneath Longmu Co Lake in western QTP.
4.3. Megathermal in the Middle Holocene (∼8500–7000 to ∼4000–3000 years B.P.)
 The middle Holocene was a period of climatic optimum, and often is called the Megathermal or Hypsithermal. The climates became increasingly drier with a dwindling intensity in wetness fluctuations and an alpine steppe had been established on the northern QTP, at least since the arid period ∼5500–4300 years B.P. [Avouac et al., 1996; X. P. Yang et al., 2002, 2004]. However, the majority of 14C ages of the thick layers of peat and humus fall within this period, indicating a warm and humid climate. The age of humic soils is 7530 ± 300 years B.P. at the depth of 4.4 m in Borehole 8 in Xidatan and the age of the ash-like carbon clayey sands on the first terrace in Nachitai (Naij Tal) is 4910 ± 100 years B.P. Many sites with remnants of human use of fire from Nachitai to Xidatan along the Kunlun He River suggest a more-clement-than-today climate and an environment conducive for human existence. The thick humus layer at HMSS 109 south of the Tanggula Shan Mountain Pass is 14C-dated at 5058 ± 443 years B.P. and that at HMSS 120 is 4470 ± 102 years B.P. The ending of the continuous deposition of thick peat layers at Qinongga and Wuma Qü River area are 3050 ± 120 and 3575 ± 80 years B.P., respectively. All those deposits were during the hypsithermal, which indirectly suggests that permafrost had largely retreated and that vast areas north of the Kunlun Shan Mountains and south of the Tanggula Shan Mountains were clear of permafrost.
 In the interior, permafrost had been subjected to persistent downward thawing to depths of ∼14–16 m and warming to even greater depths, resulting in a talik layer disconnecting the underlying permafrost from seasonal frost. Thick-layered ice was formed at depths of ∼14–16 m, which were at the relatively stable permafrost table. Because of the thawing of the shallow layer permafrost and melting of the ground ice, thermokarst lakes, lowlands and ice wedge casts were widespread on the high plateaus. Permafrost survived in islands or was deeply buried at various locations on the high plateaus. However, permafrost still existed in large expanses in the mountainous areas such as in the Kunlun Shan, Fenghuo Shan and Tanggula Shan Mountains in the interior QTP. The degradation of permafrost in the eastern area was more dramatic than in the interior and in the west. Permafrost completely vanished below 4200 m along the QKH and the permafrost tables could have been lowered to depths of ∼15–25 m at elevations from ∼4200–4400 m from Huashixia to Qingshui'he. Permafrost was retreating laterally and degrading downward, resulting in 10- to 20-m-thick residual permafrost in some areas (Figures 3 and 7). Permafrost survived in alpine areas of the eastern QTP such as in the Bayankala (Bayan Har) Shan and Animaqin (A'nyêmaqên) Shan Mountains, in the form of island permafrost.
 According to the manifestations of relict permafrost, the lower limit of permafrost (LLP) during this period was ∼300–400 m higher than today. It also is deduced that the MAATs were ∼2–3°C higher. This warmer period was persistent, which caused extensive degradation of permafrost. It is estimated that the total areas of permafrost then were about 40–50% of those at present. However, even during the middle Holocene Megathermal (about 5500 years B.P.), there were sharp cooling periods (about 6–7°C) in the sources of the Yellow River [Cheng et al., 2006].
4.4. Neoglaciation Cold Period in the Late Holocene (∼4000–3000 to 1000 years B.P.)
 With fluctuating climatic changes, the Neoglaciation returned to the cold climates of the late Holocene (∼4000–3000 to 1000 years B.P.). Mountain glaciers advanced extensively as evidenced by the three and four terminal, and lateral moraines between the terminal moraines of today and those of the Last Glaciation.
 There are a wide variety of time spans for the duration of the Neoglaciation on the QTP in the late Holocene. In the Dunde Ice Cap in the Qilian Shan Mountains on the northern QTP, the coldest periods occurred at 4000 years B.P. and 2800–2700 years B.P., with a warm interval at about 2900 years B.P. In the Cunce Ice Cap in the western Kunlun Shan Mountains, the older lateral moraine of the Neoglaciation was dated between (3983 ± 120) and (3522 ± 117) years B.P. [Zheng, 1991]. Many places on the QTP had glacier advances, such as in the southeastern QTP, including Xuedang (2980 ± 150 years B.P.) and Ruoguo (1920 ± 110 years B.P.); in the bordering eastern mountains, including the Hailuogou Glacier (3080 ± 80, 2480 ± 80, and 1550 ± 70 years B.P.) on the eastern slope of the Gonga Shan Mountains [Zheng, 1997]; and in the bordering northern mountains, including the Lenglong Ling Mountains on the northern slope of the Qilian Shan Mountains (∼2530–3110 years B.P.) [Wu, 1984].
 During this period, a permafrost layer as thick as 20 m was formed at the Borehole 8 location near Xidatan. A series of large pingos also was formed at elevations from ∼4250 to 4300 m in eastern Xidatan. Intensive cryoturbations developed in the soils on the first terrace (∼3700–3800 m in elevation) of the Kunlun He River near Nachitai (Naij Tal). Thick layer humic soils near the HMSS 109 south of the Tanggula Shan Mountain Pass and HMSS 120 also are epigenetic permafrost. The pingos, aged 2925 ± 175 years B.P., 40 km east of Shiqü County Town, Sichuan Province and the pingo group, aged 3925 ± 175 years B.P., north of K65 along the Maqin to Changma'he Highway were formed during this period. Large polygons, solifluctions and stone circles developed on the northern slopes of the Riyue Shan (3450 m), Heka'nan'shan (3600 m), and Ela (Ngöla) Shan (∼4000–4100 m) Mountains. The soils below the solifluction lobes at 4050 m developed later than the silts and sands containing humus, aged at 4395 ± 215 years B.P. The widespread dated relict permafrost and periglacial phenomena are indications of a colder climate during that period.
 By comparison with the present and earlier distribution of periglacial phenomena, it is concluded that the LLP then was about 300 m lower than today, and MAATs were about 2°C colder. Permafrost expanded after its retreat during the Megathermal period, and reached its maximum distribution by the end of the cold period about 1000 years ago. The total area of permafrost was ∼20–30% greater than today. The LLP then was at elevations of ∼3700–3800 m west of Nachitai in the north, and at ∼4400–4500 m in the Dangxiong (Damxung) Valley in the south. In the central QTP, from the Kunlun Shan to the Tanggula Shan Mountains, epigenetic permafrost as thick as 30 m developed above 4500 m on the high plateau [Ding and Guo, 1982]. This layer of permafrost could have been connected with the first layer of residual permafrost from the Megathermal. Deeply buried single- or double-layer permafrost has not yet been discovered in the central QTP.
 On the contrary, deeply buried vertical taliks, or double-layer permafrost, have been found in many areas from Huashixia to Qingshui'he along the Qinghai-Kang (western Sichuan) Highway (QKH), and in other areas to the east. Along the margins of the QTP, permafrost thawed downward to depths of ∼15–20 m during the Megathermal, but the more recent permafrost, formed during the later cold period, is less than 15 m in thicknesses (Figure 3). It also is possible that a minor warming occurred after the cold period and thawed some of the upper permafrost, leaving the deeply buried permafrost and the vertical taliks revealed in the boreholes in some areas.
4.5. Warm Period in the Late Holocene (1000–500 years B.P.)
 There were several small-scale climatic fluctuations during the late Holocene. The period 1000–500 years B.P. in China and the medieval period 1200–500 years B.P. in Europe had significant increases in temperatures that lasted for hundreds of years. The effect of warming would have impacted the thermal regimes of permafrost as deep as ∼130–150 m [Brewer, 1958], and would have thawed shallower permafrost. Because it happened recently, the remains of periglacial phenomena and permafrost have been well preserved, including the remains of collapsed pingos at Xidatan, 40 km east of Shiqu County, and at K65 along the Maqin to Changma'he Highway. The humus in the center of the collapsed pingos was formed during ∼720–625 years B.P.; and thick layers of humus were developed about 780 ± 131 years B.P. in the permafrost islands near HMSS 121 along the QTH. The lower layer of permafrost at depths of ∼9.7–12.3 m in Borehole CK1 in Di'na'rantan northeast of Huashixia and at depths of ∼11.6–15.2 m in Borehole ZK8 in Changma'he, should be the residual permafrost resulting from the downward thawing (Figure 3). Correspondingly, the top of the upper layer of relict permafrost was identified at depths of ∼7.5–9.0 m on the high plateaus and in the Fenghuo Shan Mountains. Two layers of relict permafrost can be clearly identified in Borehole CK 224 on the Chuma'er'he (Cumar'he) High Plateau. The upper permafrost table is at a depth of 8.35 m, and should belong to the warm period in the late Holocene. The lower is at the depth of 16 m, and probably remains from the Megathermal. Thick-layered ice increased in the vicinity of the permafrost tables. During the warm period of the late Holocene, the LLP and MAATs were ∼200–300 m and 1.5–2.0°C warmer than at present, respectively. The retreating of permafrost resulted in a total permafrost area of ∼20–30% less than at present.
 Some small- and medium-size lakes dried up, or became salt lakes. For example, the age of the silty soils at a depth of 0.6 m below the salt layer is 1094 ± 344 years B.P. beneath the salt lake behind HMSS 69 along the QTH; the age of the silty soils at a depth of 0.5 m below the salt layer is 1080 ± 260 in Hajiang Salt Lake east of Eling (Ngöring) Lake in a source of the Huang He (Yellow) River. The once fixed, or semifixed, sand dunes were again covered by aeolian sands, and desertification resumed.
4.6. Little Ice Age (LIA) in the Later Holocene (∼500–100 years B.P.)
 The LIA on the QTP occurred during ∼500–100 years B.P.: the climate was drying and cooling. Lakes shrunk rapidly and extensively. During the LIA, there were three cold periods (A.D. ∼1430–1540, ∼1647–1733, and ∼1778–1870), and three warmer periods (A.D. ∼1397–1429, ∼1541–1646, and ∼1734–1777) [Kang, 1996a, 1996b]. However, the cold periods lasted longer than the warmer periods, and the transitions from cold to warmer climates were sharper. Tree rings, ice cores, the glacial geology, and recent meteorological measurements support these observations [Yao, 1992; Kang, 1996a, 1996b].
 Permafrost again started to develop, with increasing thickness and areal extent. A new layer of epigenetic permafrost about 10 m thick was formed by the downward freezing of soils, connecting with the upper layer of relict permafrost remaining from the warm period during the late Holocene. Permafrost islands remained and expanded during the LIA. In some areas on the eastern QTP, the permafrost that formed during the LIA was thin, with depths of ∼1.5–8.0 m in Borehole ZK6 along the northern shore of the Eling (Ngöring) Lake and at depths of ∼5.3–8.2 m in the Qingshui'he water well along the QKH. This permafrost did not connect with the underlying relict permafrost (Figure 3).
 The temperatures during the LIA were coldest during the 17th century. The lower limit of relict rock fields along the highways in Dari (Darlag) County, Qinghai Province is 4130 m, with blocks of humus 425 ± 85 years B.P. in age underlying some of the rocks. The age of relict stone fields should be later in the LIA. At present, the LLP is higher than 4300 m, but should have been at least ∼150–200 m lower during the LIA, the MAATs about 1°C colder [Kang, 1996a], and the total area of permafrost 10% greater.
4.7. Recent Warming Period During the Last Century (∼100–0 years B.P.)
 During the past century, the climatic conditions on the QTP can be divided into three periods: (1) a warm period during ∼1871–1916, (2) a cold period during ∼1917–1968, and (3) a strong and persistent warming after 1968 [Kang, 1996a, 1996b]. Climatic warming has been apparent, especially during the past 40 years. While the global average increase in MAATs has been ∼0.3–0.6°C, since 1880, that on the QTP has been ∼0.3–0.5°C during the past 40 years, resulting in a significant retreat of permafrost [Wang, 1993; Kang, 1996b; Jin et al., 1999, 2000; Wang et al., 2000]. The depth of seasonal frost penetration has declined by ∼5–20 cm in nonpermafrost or in deep relict permafrost areas. The depth of seasonal thaw penetration has increased by ∼25–60 cm in the shallow permafrost areas. The MAGTs of permafrost have risen by ∼0.1–0.4°C over extensive areas. Permafrost has been shrinking toward the interior, with a ∼6–8% reduction of the permafrost area. The degradation in the periphery permafrost zones is the most striking, it has shrunk about 12% at Xidatan near the northern lower limit of permafrost (LLP) [Wang, 1993]; and about 20% in the Liangdao He River Basin near the southern LLP [Nan et al., 2003]. The LLP has risen 40 to 80 m.
 The vertical degradation of permafrost has disconnected seasonally frozen ground from permafrost, in many areas, forming taliks, or interannually thawed layers [Jin et al., 1999, 2000, 2006]. The thawed layer(s) have increased tens of centimeters, and can be as thick as 3 m, beneath the seasonally frozen layer. The interannually thawed layer is 2.7 to 3.5 m thick in the Qingshui'he water well along the Qinghai-Kang (western Sichuan Province) (QKH); and vertical taliks beneath the seasonally frozen ground are as thick 8 to 14 m in Borehole ZK2 at Qingshui'he along the QKH, in Borehole ZK8 at Changma'he, and at Borehole CK1 at Di'na'rantan northeast of Huashixia [Jin et al., 2006]. The taliks resulting from deep thawing of permafrost during the late Holocene were thick, and the upper layer of permafrost that formed during the LIA was thin.
5. Conclusions and Recommendations
 Since the end of the late Pleistocene, the Plateau environments have changed greatly, with a major warming occurring during 8500–3000 years B.P. Plateau permafrost responded with a lag time and damped amplitudes of change at depths, as well as with strong regional variations. Three ground ice layers with tops of relict permafrost at 2–4, 7–8, and 14–16 m have been identified by crossexamining the segregated ice, and the soluble salt contents. During the 1960s and 1970s, the new and old permafrost layers were believed to have been connected vertically. Later explorations revealed the existence of double-layer permafrost on the eastern QTP, but no evidence has been reported for the occurrence of these phenomena west of the QTH, indicating strong east-to-west variations in permafrost evolution. In Figure 7, double-layer or deep permafrost asymmetrically vertically encircles connected permafrost at roughly the −6°C MAAT isotherm. The present LLP on the QTP roughly coincides with the −3 to −2°C MAAT isotherms. It is believed that the residual permafrost with −4 to 0°C MAAT isotherms also developed during the warmer periods of the Holocene.
 A similar situation has been discovered in northeastern China [Jin et al., 2007], but not in the mountainous areas of western China. In western Siberia and European Russia, deeply buried (50–100 m in depths) permafrost continued to exist during the Megathermal, when the southern limit of vertically joined permafrost retreated to 66–67°N, but that of residual permafrost could have existed as far south as 57°–56°N in western Siberia and 59°–60°N in Europe [Baulin, 1981].
 The degradation of permafrost has happened and is happening along the QKH and in the eastern and southern Plateau. The degradation has been moving westward and northward because of the strengthening influence of the monsoons under a warming climate [Jin et al., 2006]. Therefore careful studies in the eastern QTP could be helpful in predicting the changes to permafrost in the interior and to the west. However, it also is possible that the sharp and strong indications of permafrost retreating on the eastern Plateau could be the superimposed effects of climatic warming accompanied by increasing precipitation. If that were the case, the degradation in the interior and western QTP could be much slower and might not form the deeply buried single or double-layer permafrost. The authors believe that other parts of the QTP and the central Asian mountains should have residual permafrost, although drilling has been inadequate to reveal such existence. More detailed exploration on the western plateau, especially along the edges, would be most helpful in resolving this question. If the glaciations and the evolution of permafrost can be combined, it might result in different scenarios for the evolution of permafrost with or without thick glacier coverage.
5.1. Relationships Between the Latitudinal and Elevational Permafrost During the LGM
 During the LGM (35 to 11 ka B.P.), a large portion of northern China and central Asia were under the strong influence of aeolian action and the Penultimate Glaciation, with arid and semiarid steppes and grasslands [Zhou et al., 2000; X. P. Yang et al., 2004]. The Eurasian permafrost expanded southward to 36°30′N in eastern China, to 39°N in northeastern China and to 40°30′N in western China, and connected with elevational permafrost in central Asia and on the QTP with LLPs at 1950–2950 m [Cui, 1984, 2004; Zhou et al., 2000]. However, the relict lower limit of permafrost (LLP) should have existed at 1800–1900 m. Permafrost should have been absent in the Tenggeli Desert because of its elevations of 1000–1500 m, but present in the Maowusu (Mu Us) Desert on the nearby Eerduosi (Erdos) Plateau with similar latitudes and elevations. In the Zhunge'er (Zhunger) Basin, permafrost should probably have been absent at elevations of less than 1000 m with relict LLPs present at 1400–1500 m. In northeastern China, the Sunhua Jiang-Liao He (Song-Liao) Rivers Plain, with similar elevations and latitudes, should at least have had island permafrost. These discrepancies need further evidence for interpretation. There have been vigorous discussions concerning an unified ice sheet or limited plateau ice sheets during the Quaternary [Kuhle, 1988; Shi et al., 1992b, 1995; Shi, 2002; Zheng et al., 2002], and it is possible that permafrost could have been absent beneath a kilometer or two of an extensive or limited ice sheet. With such great extents of permafrost and glaciers in the middle latitudes of central Asia, their feedback and response to climatic changes through surface processes need further field investigations and simulations at various spatiotemporal scales.
5.2. Relationships Between Aeolian Action (Desertification) and the Evolution of Permafrost
 There have been speculations and hypotheses that climatic warming and subsequently the degradation of permafrost caused enhanced aeolian erosion and widespread desertification [Wang et al., 2002]. These are based on the thought that deepening active layers and the disappearance of permafrost islands in the marginal areas would have lowered the suprapermafrost water tables, thus removing moisture available for plant growth and the resulting gradual loss of vegetation coverage. These can facilitate the saltation of fine sands and silts at the ground surface because of enhanced evaporation and periglacial wind erosion. Aeolian action intensification, as evidenced by land desertification, has occurred more or less simultaneously with permafrost expansion on the QTP since the late Pleistocene, particularly during the LGM. The periglacial environments were accompanied by frost weathering of rocks, increased aridity, and evaporation, as well as barrenness of lands immediately after the glaciations [Williams, 1994]. These combined to provide the extraordinary quantities of sand deposited in the basins from the rapid erosion of the much more strongly uplifted mountains and plateaus [Williams, 1994]. Further complications may have occurred as sediments were released by glacial meltwaters, as the snow cover declined during the Late Glaciation, and as the winds moved large quantities of sediments [Wang and Dong, 1994]. These periods during the Pleistocene provided most of the existing aeolian landforms, although few earlier aeolian landforms have survived because of their vulnerability to wind and water erosion. However, many inactive ventifacts are found in ancient periglacial environments. The QTP dune development seems to have been a feature mostly of earlier postglacial times, fostered by extensive areas of silty and sandy periglacial deposits such as glacial outwash, high winds and by the slow recolonization of vegetation. There also are interpretations that deserts, moving dunes and the recent desertification have been caused, at least partially, by uncontrolled cattle grazing [Miehe and Miehe, 2000; P. Yang et al., 2004]. However, the deserts have been there for a long time, but their areal extent probably has been on a general decline since the middle Holocene. A large portion of the QTP is, by definition, one of the driest alpine deserts on earth [Miehe et al., 2002].
5.3. Methodologies in the Study of Permafrost Evolution
 The evolution of permafrost during and since the late Pleistocene was closely related to glaciations and deglaciations, aeolian action, changes in the extents of lake/wetlands, neotectonics, regional climates and monsoons, and, to a lesser degree, to the surface coverage by vegetation, sand, snow and ice. It is difficult to precisely illustrate the evolution of permafrost without reliable evidence from these other disciplines, and the uncertainties of all paleoclimatic and paleoenvironmental research remain large [Kuhle, 1988; Shi et al., 1992b; Frenzel, 1994]. For example, the age of the uplift of the QTP is still a matter of controversy. Fort  concluded that alpine elevations were reached in pre-Quaternary times, while Xu  states that only in late Quaternary did the uplift reach elevations above 4000 m. As a result, permafrost evolution during the Pleistocene is still pending further evidence, but the evolution during the Holocene is more detailed and certain.
 The historical trends of permafrost changes are complicated because they are affected by climatic changes at various spatiotemporal scales, and they interact with other geographical entities such as glaciers, topography, rivers, lakes and deserts. The impacts from climatic changes have strong regional variations in the depths of soils encompassed, amplitudes of temperature changes and the areal extent. Therefore, even during the same period, in different areas and at different depths, degradation, formation and changing temperatures of permafrost can be occurring. Frozen soils, particularly at shallow depths, have been through many cycles of freezing and thawing, and the shifts between permafrost and talik formation during several earlier periods are complicated. In addition, the ages of permafrost-defining temperature conditions are difficult to accurately determine because the response of permafrost to climatic change lags in time with increasing depths. These all impose certain problems in reconstructing the evolutionary history of permafrost. Therefore, in order to fully explain the evolution of permafrost and periglacial environments and to apply the results in engineering designs and environmental management on the QTP, more widespread observations of permafrost changes should be pursued, especially those associated with changing climatic conditions.
5.4. Differential Distribution of Deeply Buried, or Double-Layer Relict Permafrost in the Eastern and Interior QTP
 The cold period in the late Pleistocene caused a new layer of permafrost to form. A warming trend followed in the early Holocene and greatly accelerated during the Megathermal (8500–4000 years B.P.), resulting in an extensive downward degradation of this permafrost layer. The warming was followed by a cooling and the formation of a new layer of permafrost during the Neoglaciation period (4000–1000 years B.P.). The depths of permafrost degradation, the thicknesses of the newly formed permafrost, and whether the older and newer permafrost layers were connected varied because of the regional differences in climatic changes (temperatures and precipitation).
 The regions north of the Tanggula Shan Mountains and west of the QTH, with elevations generally above 4500 to 4700 m and present MAATs less than −0.6°C, was warmed by less than 2°C during the Megathermal. No double- or single-layer permafrost has so far been discovered. It is deduced that the downward thawing of permafrost was less than 6–15 m in thickness during the Megathermal and that it refroze during Neoglaciation period and was connected to the late Pleistocene permafrost.
 In the regions south of the Tanggula Shan Mountains and east of the QTH, particularly on the northern Tibetan Plateau and in the sources of the Huang He (Yellow) River, the present MAATs range from −2 to −4°C as elevations vary from 4000 to 4500 m. These regions are on the southern and southeastern margins of the QTP, with lower latitudes and elevations. They may have been subjected to stronger southeastern monsoons during the Megathermal. It is estimated that climatic warming might have exceeded 2°C during the Megathermal. As a result, most of these regions were in a positive MAAT zone. Permafrost may have degraded strongly and extensively, resulting in deeply buried residual permafrost from 15 to 20 m downward. During the cold period in the Neoglaciation, a 10-m-layer of permafrost was again formed in many places, resulting in double-layer buried permafrost. In some places, the upper layer may have been thawed completely, resulting in a single layer of buried LGM permafrost. Unfortunately, reliable direct and absolute ages of these two layers of ground ice, and reliable interpretations of the δ18O, are still unavailable. All the dating data are the combined results of the dating of organic soils and relative dating of soil strata closely related with the ground ice layers. Further investigations of isotopic compositions and dating of the thick-layered ground ice may lead to further refinements.
 This paper was supported by the Chinese Academy of Sciences 100 Talents Project (2004) “Stability of linear engineering foundations in warm permafrost regions under a changing climate” and “Environmental changes on the Qinghai-Tibetan Plateau during the Holocene and their relationships with ecosystems” (KZCX3-SW-339-3). Max C. Brewer with the U.S. Geological Survey provided generous assistance in editing. Qihao Yu, Lanzhi Lü, and Sizhong Yang with the State Key Laboratory of Frozen Soils Engineering; Youhua Ran with the Laboratory of Geographic Information Sciences of the Cold Regions Environmental and Engineering Research Institute; Baoshan He with the Chinese Academy of Sciences Key Laboratory of Deserts and Desertification provided assistance in preparing some tables, photographs, and figures. Editor Robert Anderson, Associate Editor Tingjun Zhang, and two unidentified reviewers spent generous time working for improving the quality of the paper and have provided great insights which significantly benefited the authors. Sincere gratitude is thus acknowledged.