Arctic climate variability prior to 100 years BP

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Published: February 9, 2010, 2:49 pm
Updated: May 7, 2012, 11:18 am
Author: International Arctic Science Committee
Topic Editor: Howard Hanson
Topics:

This is Section 2.7 of the Arctic Climate Impact Assessment Lead Author: Gordon McBean. Contributing Authors: Genrikh Alekseev, Deliang Chen, Eirik Førland, John Fyfe, Pavel Y. Groisman, Roger King, Humfrey Melling, Russell Vose, Paul H.Whitfield

This section examines the record of past climate change in the Arctic with the objective of providing a context for evaluating evidence of more recent climate change and the possible impacts of future climate change. This review focuses on the past two million years (approximately), and particularly the past 20,000 years. Over geological time periods, the earth’s natural climate system has been forced or driven by a relatively small number of external factors.Tectonic processes acting very slowly over millions of years have affected the location and topography of the continents through plate tectonics and ocean spreading. Changes in the orbit of the earth occur over tens to hundreds of thousands of years and alter the amount of solar radiation received at the surface by season and by latitude. These orbital changes drive climate responses, including the growth and decay of ice sheets at high latitudes. Finally, changes in the emissivity of the sun have taken place over billions of years, but there have been shorter-term variations that occurred over decades, centuries, and millennia. As a consequence of changes in these external forces, global climate has experienced variability and change over a variety of timescales, ranging from decades to millions of years.

The sparseness of instrumental climate records prior to the 20th century (especially prior to the mid-19th century) means that estimates of climate variability in the Arctic during past centuries must rely upon indirect “proxy” paleoclimate indicators, which have the potential to provide evidence for prior large-scale climatic changes.Typically, the interpretation of proxy climate records is complicated by the presence of “noise” in which climate information is immersed and by a variety of possible distortions of the underlying climate information[1]. Careful calibration and cross-validation procedures are necessary to establish a reliable relationship between a proxy indicator and the climatic variable or variables of interest, providing a “transfer” function that allows past climatic conditions to be estimated. Of crucial importance is the ability to date different proxy records accurately in order to determine whether events occurred simultaneously, or whether some events lagged behind others.

Sources of paleoclimatic information in the Arctic are limited to a few, often equivocal types of records, most of which are interpreted as proxies for summer temperature. Little can be said about winter paleoclimate[2]. Only ice cores, tree rings, and lake sediments provide continuous high-resolution records. Coarsely resolved climate trends over several centuries are evident in many parts of the Arctic, including:

  • the presence or absence of ice shelves deduced from driftwood frequency and peat growth episodes and pollen content;
  • the timing of deglaciation and maximum uplift rates deduced from glacio-isostatic evidence as well as glacial deposits and organic materials over-ridden by glacial advances or exposed by ice recession;
  • changes in the range of plant and animal species to locations beyond those of today[3]; and
  • past temperature changes deduced from the geothermal information provided by boreholes.

In contrast, large-scale continuous records of decadal, annual, or seasonal climate variations in past centuries must rely upon sources that resolve annual or seasonal climatic variations. Such proxy information includes tree-ring width and density measurements; pollen, diatom, and sediment changes from laminated sediment cores; isotopes, chemistry, melt-layer stratigraphy, acidity, pollen content, and ice accumulation from annually resolved ice cores; and the sparse historical documentary evidence available for the past few centuries.

Information from individual paleoclimate proxies is often difficult to interpret and multi-proxy analysis is being used increasingly in climate reconstructions. Taken as a whole, proxy climate data can provide global- scale sampling of climate variations several centuries into the past, and have the potential to resolve both large-scale and regional patterns of climate change prior to the instrumental period.

Pre-Quaternary Period (2.7.1)

Over the course of millions of years, the Arctic has experienced climatic conditions that have ranged from one extreme to the other. Based on the fossil record, 120 to 90 million years before present (My BP) during the mid-Cretaceous Period, the Arctic was significantly warmer than at present, such that arctic geography, atmospheric composition, ocean currents, and other factors were quite different from at present. In contrast, as recently as 20 ky (thousand years) BP, the Arctic was in the grip of intense cold and continental-scale glaciation was at its height.This was the latest of a series of major glacial events, which, together with intervening warm periods, have characterized the past two million years of arctic environmental history.

The most recent large-scale development and build-up of ice sheets in the Arctic probably commenced during the Late Tertiary Period (38 to 1.6 My BP). Ice accumulation at that time may have been facilitated initially by plate tectonics[4]. According to Maslin et al., the onset of Northern Hemisphere glaciation began with a significant build-up of ice in southern Greenland. However, progressive intensification of glaciation does not seem to have begun until 3.5 to 3 My BP, when the Greenland Ice Sheetexpanded to include northern Greenland. Maslin et al.[5] suggested that the Eurasian Arctic, northeast Asia, and Alaska were glaciated by about 2.7 My BP and northeast America by 2.5 My BP. Maslin et al.[6] suggested that tectonic changes, including the deepening of the Bering Strait, were too gradual to be responsible for the speed of Northern Hemisphere glaciation, but suggested that tectonic changes brought the global climate to a critical threshold while relatively rapid variations in the orbital parameters of the earth triggered the glaciation.

Quaternary Period (2.7.2)

The Quaternary Period (the last 1.6 My) has been characterized by periodic climatic variations during which the global climate system has switched between interglacial and glacial stages, with further subdivision into stadials (shorter cold periods) and interstadials (shorter mild episodes). Glacial stages are normally defined as cold phases with major glacier and ice sheet expansion. Interglacials are defined as warm periods when temperatures were at least as high as during the present Holocene interglacial.

This interglacial–glacial–interglacial climate oscillation has been recurring with a similar periodicity for most of the Quaternary Period, although each individual cycle appears to have had its own idiosyncrasies in terms of the timing and magnitude of specific events. It has been estimated that there have been between 30 and 50 glacial/interglacial cycles during the Quaternary Period [7], primarily driven by changes in the orbit of the earth[8].

One of the earliest (and certainly the most well known) hypotheses concerning the effects of orbital configuration on glacial cycles is described by Milankovitch[9], who presents the argument that a decrease in summer insolation is critical for glacial initiation. Low summer insolation occurs when the tilt of the axis of rotation of the earth is small; the poles are pointing less directly at the sun; the Northern Hemisphere summer solstice is farthest from the sun; and the earth’s orbit is highly eccentric.The key orbital parameters involved include changes in the eccentricity of the orbit of the earth with a period of 100 ky; the tilt of the axis of rotation of the earth, which oscillates between 22.2º and 24.5º with a period of 41 ky; and the position of the earth within its elliptical orbit during the Northern Hemisphere summer, which changes over a period of 23 ky. Bradley[10] pointed out that such orbitally induced radiation anomalies are especially significant at high latitudes in summer when daylight persists for 24 hours.The latitude most sensitive to low insolation values is 65º N, the latitude at which most ice sheets first formed and lasted longest during the last glaciation. The amount of summer insolation reaching the top of the atmosphere at 65º N and nearby latitudes can vary by as much as ±12% around the long-term mean value. Changes in winter insolation also occur with exactly the opposite timing, but they are not considered important to ice-sheet survival.When summer insolation is weak, less radiation is delivered to the surface at high latitudes, resulting in lower regional temperatures. Lower temperatures reduce the summer ablation of the ice sheets and allow snow to accumulate and ice sheets to grow. Once ice sheets are created, they contribute to their own positive mass balance by growing in elevation, reaching altitudes of several kilometers where prevailing temperatures favor the accumulation of snow and ice.

Other hypotheses concerning the causes of glacial initiation include that of Young and Bradley[11], who argued that the meridional insolation gradient is a critical factor for the growth and decay of ice sheets through its control over poleward moisture fluxes during summer and resultant snowfall. Snowfall could increase in a cooler high-latitude climate through enhanced storm activity forced by a greater latitudinal temperature gradient. Ruddiman and McIntyre[12] and Miller and de Vernal[13] suggested that the North Atlantic Ocean circulation is important for modulating ice volume. There is evidence that the North Atlantic remained warm during periods of ice growth and they proposed that enhanced meridional temperature gradients and evaporation rates during the winter enhanced snow delivery to the nascent ice sheets of northeastern Canada. An active thermohaline circulation in the North Atlantic could increase the moisture supply to sites of glacial initiation through its ability to maintain warmer sea surface temperatures, limit sea-ice growth, and allow for greater evaporation from the ocean surface.

These external forcing mechanisms in turn cause responses and chain reactions in the internal elements of the earth’s system[14]. Changes in one internal element of the system can cause responses in other elements.These can lead to feedback effects that can amplify or attenuate the original signal.Thus, ice sheets, ice caps, and glaciers play an important role in the global climate system. Glacial advance and retreat may therefore be both a consequence and a cause of climate change[15].

Given the inherent errors in dating techniques, gaps in the stratigraphic record, and the varying rates of response of different biological proxy indicators, there is considerable uncertainty about the timing of specific events and whether climate changes were truly synchronous in different regions.The errors and uncertainties tend to be amplified farther back in the paleoclimatic record, particularly in the Arctic, where much of the paleoclimatic evidence from earlier parts of the Quaternary Period has been removed or obfuscated as a result of later glaciations. Consequently, this review of climate variability in the Arctic during the Quaternary Period focuses first on the more complete and reliable evidence of climate conditions immediately prior to the onset of the most recent glacial–interglacial oscillation (~130 ky BP).This is followed by a brief review of conditions during the Last Glacial Maximum (~20 ky BP) and the subsequent period of deglaciation, and culminates in a review of the evidence of climatic variability during the present interglacial, the Holocene.

Last interglacial and glaciation (2.7.3)

Last interglacial:The Eemian (2.7.3.1)

Climatic conditions during interglacial periods are generally considered to be broadly comparable to present-day conditions. The most recent interglacial, the Eemian, extended from the end of the penultimate glaciation about 130 ky BP until about 107 ky BP when the last glacial period began[16]. The Eemian is often regarded as a typical interglacial event with characteristics including relatively high sea level, a retreat to minimum size of global ice sheets, and the establishment of biotic assemblages that closely parallel those at present. According to most proxy data, the last interglacial was slightly warmer everywhere than at present[17]. Brigham-Grette and Hopkins[18] reported that during the Eemian the winter sea-ice limit in Bering Strait was at least 800 km farther north than today, and that during some summers the Arctic Ocean may have been ice-free. The northern treeline was more than 600 km farther north, displacing tundra across all of Chukotka[19]. Western European lake pollen records show deciduous forests (characteristic of warmer conditions) across much of Western Europe that were abruptly replaced by steppic taxa characteristic of colder conditions; this shift is associated with a cold event at 107 ky BP. This relatively prolonged warm period is also detected in northeast Atlantic marine sediments, but is not evident throughout the North Atlantic. Faunal and lithic records from the polar North Atlantic[20] indicate that the first abrupt cooling occurred around 118 to 117 ky BP[21].

Evidence of warmer conditions in the Arctic than exist at present is provided by a re-evaluation of the oxygen isotope ratio (d18O) record obtained from Greenland ice core samples[22]. These authors suggest that the Greenland Ice Sheet was considerably smaller and steeper during the Eemian than at present and probably contributed 4 to 5.5 m to global sea level during that period. This implies that the climate of Greenland during the Eemian was more stable and warmer than previously thought, and the consequent melting of the Greenland Ice Sheet more significant.

Some researchers suggest that a paleoclimatic reconstruction of the Eemian provides a means of establishing the mode and tempo of natural climate variability with no anthropogenic influence. However, a general lack of precise absolute timescales and regionally-to- globally synchronous stratigraphic markers makes long-distance correlation between sites problematic, and inferred terrestrial changes are difficult to place within the temporal framework of changes in ice volume and sea level[23].

Last glaciation: Wisconsinan/Weichselian (2.7.3.2)

Although the timing of the end of the Eemian interglacial is subject to some uncertainty, high-resolution North Atlantic marine sediment records indicate that the Eemian ended with abrupt changes in deep-water flow occurring over a period of less than 400 years[24]. Evidence in marine sediments of an invasion of cold, low salinity water in the Norwegian Sea at this time has been linked by Cortijo et al.[25] to a reduction in warm-water transport by the North Atlantic Drift and the thermohaline circulation. It seems likely that this contributed to the onset of widespread arctic glaciation in sensitive areas.

Following the initial cooling event (~107 ky BP), climatic conditions often changed suddenly, followed by several thousand years of relatively stable climate or even a temporary reversal to warmth. Overall, however, there was a decline in global temperatures.The boundaries of the boreal forests retreated southward and fragmented as conditions grew colder. Large ice sheets began to develop on all the continents surrounding the Arctic Ocean. The point at which the global ice extent was at its greatest (~24 to 21 ky BP) is known as the Last Glacial Maximum (LGM)[26].

At its maximum extent, the Laurentide Ice Sheet extended from the Arctic Ocean in the Canadian Archipelago to the midwestern United States in the south, and from the Canadian Cordillera to the eastern edge of the continent. Local ice sheets also covered the Alaska and Brooks Ranges in Alaska.The Eurasian and Laurentide Ice Sheets were responsible for most of the glacio-eustatic decrease in sea level (about 120 m) during the LGM.The pattern of postglacial isostatic rebound suggests that the ice was thickest over Hudson Bay.The different parts of the Laurentide Ice Sheet reached their maximum extent between 24 and 21 ky BP[27].The Innuitian ice buildup appears to have culminated in the east after 20.5 ky BP. Dyke et al.[28] suggested that the entire ice sheet system east of the Canadian Cordillera responded uniformly to changes in climate. In contrast, the Cordilleran Ice Sheet did not reach its maximum extent until 15.2 to 14.7 ky BP, well after the LGM and the insolation minimum at approximately 21 ky BP[29].This out-of-phase response may be attributable to the effects of growth of the Cordilleran Ice Sheet, which would have intercepted moisture transport to the interior plains at the expense of the Laurentide Ice Sheet. During its maximum extent, the Laurentide Ice Sheet was more than twice the size of the Eurasian Ice Sheet. Changes in climate during the LGM are discussed by the IPCC[30].

The Eurasian Ice Sheet initiation began 28 ky BP as a result of temperature changes that lowered equilibrium line altitudes across the Scandinavian mountains, Svalbard, Franz Josef Land, and Novaya Zemlya. Ice flow north from Scandinavia and south from Svalbard, in conjunction with eustatic sea-level decreases, caused the margin of the ice sheet to migrate into the Barents Sea. Complete glaciation of the Barents Sea by a grounded ice sheet was achieved by 20 ky BP.The ice sheet at its maximum extent covered Scandinavia and the Barents Shelf and included a marine-based margin along the northern Barents Shelf, the western Barents Sea, western Scandinavia, and northern Great Britain and Ireland. The eastern margin of the ice sheet is generally thought to have been located west of the Taymir Peninsula[31]. It appears that at the LGM, cold arid conditions persisted across much of eastern Siberia and that an East Siberian Sea Ice Sheet did not exist[32], as some have claimed[33]. Glaciers in the Urals at the LGM appear to have been confined to the mountain valleys, rather than coalescing with the Barents-Kara Ice Sheet, as happened during previous glaciations[34].

Last glacial/interglacial transition through to mid-Holocene (2.7.4)

Last glacial/interglacial transition (2.7.4.1)

The most extreme manifestation of climate change in the geological record is the transition from full glacial to full interglacial conditions. Following the LGM (~24 to 21 ky BP), temperatures close to those of today were restored by approximately 10 ky BP[35].

The inception of warming appears to have been very rapid[36]. The rate of temperature change during the recovery phase from the LGM provides a benchmark against which to assess rates of temperature change in the late 20th century. Available data indicate an average warming rate of about 2 ºC per millennium between about 20 and 10 ky BP in Greenland, with lower rates for other regions. On the other hand, very rapid temperature increases at the start of the Bølling- Allerød period (14.5 ky BP)[37] or at the end of the Younger Dryas (~11 ky BP) may have occurred at rates as large as 10 ºC per 50 years over substantial areas of the Northern Hemisphere. Almost synchronously, major vegetation changes occurred in Europe and North America and elsewhere[38].There was also a pronounced warming of the North Atlantic and North Pacific[39].

Oxygen isotope measurements from Greenland ice cores demonstrate that a series of rapid warm and cold oscillations, called Dansgaard–Oeschger (D-O) events, punctuated the last glaciation, often taking Greenland and northwestern Europe from a full-glacial climate to conditions about as warm as at present (Fig. 2.15). For the period between 115 and 14 ky BP, 24 of these short-lived warm events have been identified in the Greenland ice core data, although many lesser warming events also occurred[40]. According to Lang et al.[41], associated temperature changes may have been as great as 16 ºC.The D-O oscillations are correlated with sea surface temperature variations derived from several North Atlantic deep-sea cores[42]. From the speed of the climate changes recorded in the Greenland Ice Sheet[43], it is widely thought that the complete change in climate occurred, at least regionally, over only a few decades. These interstadials lasted for varying periods of time, usually a few centuries to about 2000 years, before equally rapid cooling returned conditions to their previous state. The evidence from high-resolution deep sea cores from the North Atlantic[44] suggests that during at least the past 30000 years, interstadials tended to occur at the warmer points of a background North Atlantic temperature cycle that had a periodicity of approximately 1,500 years.

Heinrich events appear to be the most extreme of a series of sudden, brief cold events that seem to have occurred very frequently over the past 115 000 years, apparently tending to start at the low point of the same 1500-year temperature cycle. Heinrich events occurred during times of decreasing sea surface temperatures in the form of brief, exceptionally large discharges of icebergs in the North Atlantic from the Laurentide and European Ice Sheets that left conspicuous layers of detrital material in deep-sea sediments. Accompanying the Heinrich events were large decreases in the oxygen isotope ratio of planktonic foraminifera, providing evidence of lowered surface salinity probably caused by melting of drifting ice[45]. Heinrich events appear at the end of a series of saw-tooth-shaped temperature cycles known as Bond cycles. During the Pleistocene Epoch, each cycle was characterized by the relatively warm interstadials becoming progressively cooler.

Fig. 2.15. Temperature change over the past 100 ky (departure from present conditions) reconstructed from a Greenland ice core[46].

Deep-sea cores also show the presence of ice-rafting cycles in the intervals between Heinrich events[47].The duration of these ice-rafting cycles varies between 2000 and 3000 years and they closely coincide with the D-O events. A study of the ice-rafted material suggests that, coincident with the D-O cooling, ice within the Icelandic ice caps and within or near the Gulf of Saint Lawrence underwent nearly synchronous increases in rates of calving.The Heinrich events reflect the slower rhythm of iceberg discharges into the North Atlantic, probably from Hudson Strait.

Air temperature, sea surface temperature, and salinity variations in the North Atlantic are associated with major changes in the thermohaline circulation. A core from the margin of the Faroe-Shetland Channel covering the last glacial period reveals numerous oscillations in benthic and planktonic foraminifera, oxygen isotopes, and ice-rafted detritus[48]. These oscillations correlate with the D-O cycles, showing a close relationship between the deep-ocean circulation and the abrupt climatic changes during the last glaciation. It is increasingly apparent that large, globally linked environmental feedbacks were involved in the generation of the large, rapid temperature increases that occurred during glacial termination and the onset of D-O events.

During the last glacial–interglacial transition, the movements of the North Atlantic Polar Front have been described as pivoting around locations in the western North Atlantic. Iceland, situated in the middle of the North Atlantic, has glaciers sensitive to changes in oceanic and atmospheric frontal systems[49].The late-glacial (subsequent to the LGM) records from Iceland indicate that relatively warm Atlantic water reached Iceland during the Bølling– Allerød Interstadial, with a short cooling period corresponding with the Older Dryas. Karpuz et al.[50] suggested that the marine polar front was located close to Iceland during the Bølling–Allerød, and Sarnthein et al.[51] concluded that sea-surface circulation was mainly in Holocene interglacial mode after 12.8 ky BP. Warm episodes were also associated with higher seasurface temperatures and the presence of oceanic convection in the Norwegian Greenland Sea. Cold episodes were associated with low sea surface temperatures, low salinity, and no convection[52]. Subsequent to the initial phase of deglaciation on Spitsbergen (~13 to 12.5 ky BP), most of the eastern and northern Barents Sea was deglaciated during the Bølling–Allerød Interstadial[53]. Vorren and Laberg[54] speculate on the existence of a close correlation between summer air temperatures and the waxing and waning phases of the northern Fennoscandian and southern Barents Sea Ice Sheets. Most of the ice-sheet decay is attributed by these authors to increases in air temperature, which caused thinning of the ice sheets, making them susceptible to decoupling from the seabed and increased calving.

After a few thousand years of recovery, the Arctic was suddenly plunged back into a new and very short-lived cold event known as the Younger Dryas[55] leading to a brief advance of the ice sheets.The central Greenland ice core record (from the Greenland Ice Core Project and the Greenland Ice Sheet Project indicates that the return to the cold conditions of the Younger Dryas from the incipient interglacial warming 13 ky BP took place in less than 100 years.The warming phase at the end of the Younger Dryas, which took place about 11 ky BP and lasted about 1300 years, was very abrupt and central Greenland temperatures increased by 7 ºC or more in a few decades.

Note: In this chapter, when describing changes in arctic climate, the words possible, probable, and very probable are used to indicate the level of confidence the authors have that the change really did occur, recognizing the limitations of the observing system and paleoclimatic reconstructions of arctic climate.)

Early to mid-Holocene (2.7.4.2)

Following the sudden end of the Younger Dryas, the Arctic entered several thousand years of conditions that were warmer and probably moister than today. Peak high-latitude summer insolation (Milankovitch orbital forcing) occurred during the earliest Holocene, with a maximum radiation anomaly (approximately 8% greater than at present) attained between 10 and 9 ky BP[56]. Although most of the Arctic experienced summers that were warmer (1–2 ºC) than at present during the early to middle Holocene, there were significant spatial differences in the timing of this warm period.This was probably due to the effects of local residual land-ice cover and local sea surface temperatures[57]. In most of the ice cores from high latitudes, the warm period is seen at the beginning of the Holocene (about 11 to 10 ky BP). In contrast, central Greenland[58] and regions downwind of the Laurentide Ice Sheet, including Europe[59] and eastern North America[60], did not warm up until after 8 ky BP. The rapid climatic events of the last glacial period and early Holocene are best documented in Greenland and the North Atlantic and may not have occurred throughout the Arctic. This early Holocene warm period appears to have been punctuated by a severe cold and dry phase about 8200 years ago, which lasted for less than a century, as recorded in the central Greenland ice cores[61].

Glacio-isostatic evidence indicates deglaciation was underway by the beginning of the Holocene and maximum uplift occurred between 8 and 7 ky BP and even earlier in many areas. By 9 ky BP, Spitsbergen glaciers had retreated to or beyond their present day positions[62] and the marine faunal evidence suggest that this period was as warm if not warmer than at present along the west and north coasts of Svalbard[63].The retreat of the largest of the glaciers along the Gulf of Alaska began as early as 16 to 14 ky BP[64]. Although much of the high and mid-Canadian Arctic remained glaciated, warm summers are clearly registered by enhanced summer melting of the Agassiz Ice Cap[65]. Following the large, abrupt change in stable-isotope ratios marking the end of the last glaciation, d18O profiles from Agassiz Ice Cap cores show a rapid warming trend that reached a maximum between approximately 9 and 8 ky BP.

Deglacial marine sediments in Clements Markham Inlet on the north coast of Ellesmere Island resemble those characteristic of temperate (as opposed to polar) tidewater glaciers, suggesting that climatic conditions in the early Holocene were significantly warmer there than today[66]. Glaciers had retreated past present-day termini in some areas by 7.5 ky BP. Increasing sea surface temperatures in Baffin Bay enhanced precipitation on Baffin Island[67], leading to a widespread early Holocene glacial advance along the east coast. Marine mammals and boreal mollusks were present far north of their present-day range by 7.5 to 6.5 ky BP, as were many species of plants between 9.2 and 6.7 ky BP[68]. Caribou were able to survive in the northernmost valleys of Ellesmere Island and Peary Land by 8.5 ky BP or earlier. Such evidence indicates very warm conditions early in the Holocene (before 8 ky BP).

Early Holocene summer temperatures similar to those at present have been reconstructed in Arctic Fennoscandia by numerous studies using a range of proxies and multiproxy analyses[69]. However, abrupt climatic variations were characteristic of the early Holocene, with distinct cool episodes around 9.2, 8.6, and 8.2 ky BP[70].The most recent of these events might be connected to the widely known “8.2 ky event”, which affected terrestrial and aquatic systems in northern Fennoscandia[71]. Hantemirov and Shiyatov[72] report that open larch forests were already growing in the Yamal Peninsula of northwestern Siberia 10.5 to 9 ky BP and that the most favorable period for tree growth lasted from 9.2 to 8 ky BP. During the early Holocene, reconstructed mean temperature anomalies for the warmest month, based on pollen data across Northern Europe, show temperatures comparable to those at present[73].Temperatures then increased around 6 ky BP, with the onset of this increase delayed to around 9 ky BP in the east.

Boreal forest development across northern Russia (including Siberia) commenced by 10 ky BP[74]. Over most of Russia, forests advanced to or near the current arctic coastline between 9 and 7 ky BP, and retreated to their present position by between 4 and 3 ky BP. Forest establishment and retreat were roughly synchronous across most of northern Russia, with the exception of the Kola Peninsula, where both appear to have occurred later. During the period of maximum forest extension, the mean July temperature along the northern coastline of Russia may have been 2.5 to 7.0 ºC warmer than present.

The Arctic appears to have been relatively warm during the mid-Holocene, although records differ spatially, temporally, and by how much summer warmth they suggest (relative to the early Holocene insolation maximum)[75].

A review of the Holocene glaciation record in coastal Alaska[75] suggests that glacier fluctuations in arctic, central interior, and southern maritime Alaska were mostly synchronous. Ager[76] and Heusser[77] report that pollen records indicate a dramatic cooling about 3.5 ky BP and suggest an increase in precipitation and storminess in the Gulf of Alaska accompanied by a rejuvenation of glacial activity. In northern Iceland, the Holocene record of glacier fluctuations indicates two glacial advances between 6 and 4.8 ky BP[78].

In the Canadian Arctic, interior regions of Ellesmere Island appear to have retained extensive Innuitian and/or plateau ice cover until the mid-Holocene[79], after which ice margins retreated to positions at or behind those at present. Restricted marinemammal distributions imply more extensive summer sea ice between 8 and 5 ky BP, and hence cooler conditions[80]. However, marine conditions at 6 ky BP warmer than those at present are suggested by analyses of marine microfossils[81] performed over a broad area from the high Arctic to the Labrador Sea via Baffin Bay. A multi-proxy summary of marine and terrestrial evidence from the Baffin sector[82] suggests that warming began around 8 ky BP, intensified at 6 ky BP, and that conditions had cooled markedly by 3 ky BP.

Dugmore[83] demonstrated that, in Iceland, the Sólheimajökull glacier extended up to 5 km beyond its present limits between 7 and 4.5 ky BP. Major ice sheet advances also occurred before 3.1 ky BP and between 1.4 and 1.2 ky BP. In the 10th century (1 ky BP), this glacier was also larger than during the period from AD 1600 to 1900, when some other glaciers reached their maximum Holocene extent. Stötter et al.[84] suggested that major glacier advances in northern Iceland occurred at around 4.7, 4.2, 3.2–3.0, 2.0, 1.5, and 1.0 ky BP.

Evidence for a mid-Holocene thermal maximum in Scandinavia is considerable, and based on a wide range of proxies[85]. Treelines reached their maximum altitude (up to 300 m higher than at present)[86], and glaciers were greatly reduced or absent[87]. Pollen and macrofossil records from the Torneträsk area in northern Swedish Lapland indicate optimal conditions for Scots pine (Pinus sylvestris) from 6.3 to 4.5 ky BP[88] and records of treeline change in northern Sweden show high-elevation treelines around 6 ky BP[89]. These data indicate an extended period in the early to mid- Holocene when Scandinavian summer temperatures were 1.5 to 2 ºC higher than at present.

Tree-ring data from the Torneträsk area indicate particularly severe climatic conditions between 2.6 and 2 ky BP (600–1 BC).This period includes the greatest range in ring-width variability of the past 7400 years in this area, indicating a highly variable but generally cold climate[90]. This period is contemporary with a major glacial expansion in Scandinavia when many glaciers advanced to their Holocene maximum position[91] with major effects on human societies[92].

Especially severe conditions in northern Swedish Lapland occurred 2.3 ky BP (330 BC), with tree-ring data indicating a short-term decrease in mean summer temperature of about 3 to 4 ºC. A catastrophic drop in pine growth at that time is also reported by Eronen et al.[93], who state that this was the most unfavorable year for the growth of treeline pines in Finnish Lapland in the past 7,500 years. Reconstructed Holocene summer temperature changes in Finnish Lapland, based on proxy climate indicators in sediments from Lake Tsuolbmajavri, show an unstable early Holocene between 10 and 8 ky BP in which inferred July air temperatures were about the same as at present most of the time, but with three successive cold periods at approximately 9.2, 8.6, and 8.2 ky BP, and a “thermal maximum” between approximately 8 and 5.8 ky BP, followed by an abrupt cooling[94]. Dated subfossils (partially fossilized organisms) show that the pine treeline in northwestern Finnish Lapland retreated a distance of at most 70 km during this cooling, but that the shift was less pronounced in more easterly parts of Lapland[95].

Hantemirov and Shiyatov[96] reported that the most favorable period for tree growth in the Yamal Peninsula of northwestern Siberia lasted from 9.2 to 8 ky BP. At that time, the treeline was located at 70º N.Then, until 7.6 ky BP, temperatures decreased but this did not result in any significant shift in the treeline. The treeline then moved south until, by 7.4 ky BP, it was located at approximately 69º N. It remained here until 3.7 ky BP when it rapidly retreated (~20 km) to within 2 to 3 km north of its present position south of the Yamal Peninsula. This retreat in the space of only 50 years coincides with an abrupt and large cooling as indicated in the tree-ring data. This cooling event may have been associated with the eruption of the Thera (Santorini) volcano in the southern Aegean around 3.6 ky BP.

Tree-ring data from the Kheta-Khatanga plain region and the Moyero-Kotui plateau in the eastern part of the Taymir Peninsula indicate climatic conditions more favorable for tree growth around 6 ky BP, as confirmed by increased concentrations of the stable carbon isotope 13C in the annual tree rings[97]. The growth of larch trees at that time was 1.5 to 1.6 times greater than the average radial growth of trees during the last 2,000 years, and the northern treeline is thought to have been situated at least 150 km farther north than at present, as indicated by the presence of subfossil wood of that age in alluvial deposits of the Balakhnya River. During the past 6000 years, the eastern Taymir tree-ring chronologies show a significant and progressive decrease in tree growth and thus temperature.

Last millennium (2.7.5)

Over the last millennium, variations in climate across the Arctic and globally have continued. The term “Medieval Warm Period”, corresponding roughly to the 9th to the mid-15th centuries, is frequently used but evidence suggests that the timing and magnitude of this warm period varies considerably worldwide[98]. Current evidence does not support a globally synchronous period of anomalous warmth during that time frame, and the conventional term of “Medieval Warm Period” appears to have limited utility in describing trends in hemispheric or global mean temperature changes.

The Northern Hemisphere mean temperature estimates of Mann M. et al.[99], and Crowley and Lowery[100], show that temperatures during the 11th to the 14th centuries were about 0.2 ºC higher than those during the 15th to the 19th centuries, but somewhat below the temperatures of the mid-20th century.The longterm hemispheric trend is best described as a modest and irregular cooling from AD 1000 to around 1850 to 1900, followed by an abrupt 20th-century warming.

Regional evidence is, however, quite variable. Crowley and Lowery[101] show that western Greenland exhibited local anomalous warmth only around AD 1000 (and to a lesser extent, around AD 1400), and experienced quite cold conditions during the latter part of the 11th century. In general, the few proxy temperature records spanning the last millennium suggest that the Arctic was not anomalously warm throughout the 9th to 14th centuries[102].

In northern Swedish Lapland, Scots pine tree-ring data indicate a warm period around AD 1000 that ended about AD 1100 when a shift to a colder climate occurred[103]. In Finnish Lapland, based on a 7500-year Scots pine tree-ring record, Helama et al.[104] reported that the warmest nonoverlapping 100-year period in the record is AD 1501 to 1600, but AD 1601 was unusually cold. Other locations in Fennoscandia and Siberia were also cold in AD 1601, and Briffa et al.[105] linked the cold conditions to the AD 1600 eruption of the Huaynaputina volcano in Peru. In northern Siberia, and particularly east of Taymir where the most northerly larch forests occur, long-term temperature trends derived from tree rings indicate the occurrence of cool periods during the 13th, 16th to 17th, and early 19th centuries.The warmest periods over the last millennium in this region were between AD 950 and 1049,AD 1058 and 1157, and AD 1870 and 1979.A long period of cooling began in the 15th century and conditions remained cool until the middle of the 18th century[106].

For the most part, “medieval warmth” appears to have been restricted to areas in and around the North Atlantic, suggesting that variability in ocean circulation may have played a role. Keigwin and Pickart[107] suggested that the temperature contrasts between the North Atlantic and other areas were associated with changes in ocean currents in the North Atlantic and may to a large extent reflect centuryscale changes in the NAO.

By the middle of the 19th century, the climate of the globe and the Arctic was cooling. Overall, the period from 1550 to 1900 may have been the coldest period in the entire Holocene[108].This period is usually called the “Little Ice Age” (LIA), during which glaciers advanced on all continents.The LIA appears to have been most clearly expressed in the North Atlantic region as altered patterns of atmospheric circulation[109]. Unusually cold, dry winters in central Europe (e.g., 1 to 2 ºC below normal during the late 17th century) were very probably associated with more frequent flows of continental air from the northeast[110]. Such conditions are consistent with the negative or enhanced easterly wind phase of the NAO, which implies both warm and cold anomalies over different regions of the North Atlantic sector. Although the term LIA is used for this period, there was considerable temporal and spatial variability across the Arctic during this period.

Ice shelves in northwestern Ellesmere Island probably reached their greatest extent in the Holocene during this interval. On the Devon Island Ice Cap, 1550 to 1620 is considered to have been a period of net summer accumulation, with very extensive summer sea ice in the region.There is widespread evidence of glaciers reaching their maximum post-Wisconsinan positions during the LIA, and the lowest d18O values and melt percentages for at least 1000 years are recorded in ice cores for this interval. Mann M. et al.[111] and Jones et al.[112] supported the theory that the 15th to 19th centuries were the coldest of the millennium for the Northern Hemisphere overall. However, averaged over the Northern Hemisphere, the temperature decrease during the LIA was less than 1 ºC relative to late 20thcentury levels[113]. Cold conditions appear, however, to have been considerably more pronounced in particular regions during the LIA. Such regional variability may in part reflect accompanying changes in atmospheric circulation. Overpeck et al.[114] summarized arctic climate change over the past 400 years.

There is an abundance of evidence from the Arctic that summer temperatures have decreased over approximately the past 3500 years. In the Canadian Arctic, the melt record from the Agassiz ice core indicates a decline in summer temperatures since approximately 5.5 ky BP, especially after 2 ky BP. In Alaska, widespread glacier advances were initiated at approximately 700 ky BP and continued through the 19th century[115]. During this interval, the majority of Alaskan glaciers reached their Holocene maximum extensions.The pattern of LIA glacier advances along the Gulf of Alaska is similar on decadal timescales to that of the well-dated glacier fluctuations throughout the rest of Alaska.

There is a general consensus that throughout the Canadian Archipelago, the late Holocene has been an interval of progressive cooling (the “Neoglacial”, culminating in the LIA), followed by pronounced warming starting about 1840[116]. According to Bourgeois et al.[117], the coldest temperatures of the entire Holocene were reached approximately 100 to 300 years ago in this region. Others, working with different indicators, have suggested that Neoglacial cooling was even greater in areas to the south of the Canadian Archipelago[118]. Therefore, even if the broad pattern of Holocene climatic evolution is assumed to be coherent across the Canadian Archipelago, the available data suggest regional variation in the amplitude of temperature shifts.

The most extensive data on the behavior of Greenland glaciers apart from the Greenland Ice Sheet come from Maniitsoq (Sukkertoppen) and Disko Island. Similar to the inland ice-sheet lobes, the majority of the local glaciers reached their maximum Neoglacial extent during the 18th century, possibly as early as 1750. Glaciers started to retreat around 1850, but between 1880 and 1890 there were glacier advances. In the early 20th century, glacier recession continued, with interruptions by some periods of advance. The most rapid glacial retreat took place between the 1920s and 1940s. In Iceland, historical records indicate that Fjallsjökull and Breidamerkurjökull reached their maximum Holocene extent during the latter half of the 19th century[119]. Between 1690 and 1710, the Vatnajökull outlet glaciers advanced rapidly and then were stationary or fluctuated slightly. Around 1750 to 1760 a significant re-advance occurred, and most of the glaciers are considered to have reached their maximum LIA extent at that time[120]. During the 20th century, glaciers retreated rapidly. During the LIA, Myrdalsjökul and Eyjafjallsjökull formed one ice cap, which separated in the middle of the 20th century into two ice caps[121]. Drangajökull, a small ice cap in northwest Iceland, advanced across farmland by the end of the 17th century, and during the mid-18th century the outlet glaciers were the most extensive known since settlement of the surrounding valleys. After the mid-19th century advance, glaciers retreated significantly. On the island of Jan Mayen, some glaciers reached their maximum extent around 1850.The glaciers subsequently experienced an oscillating retreat, but with a significant expansion around 1960[122].

In northeastern Eurasia, long-term temperature trends derived from tree rings close to the northern treeline in east Taymir and northeast Yakutia indicate decreasing temperatures during the LIA[123].

Variations in arctic climate over the past 1000 years may have been the result of several forcing mechanisms. Bond et al.[124] suggested variations in solar insolation. Changes in the thermohaline circulation or modes of atmospheric variability, such as the AO, may also have been primary forcing mechanisms of century- or millennial-scale changes in the Holocene climate of the North Atlantic. It is possible that solar forcing may excite modes of atmospheric variability that, in turn, may amplify climate changes.The Arctic, through its linkage with the Nordic Seas, may be a key region where solar-induced atmospheric changes are amplified and transmitted globally through their effect on the thermohaline circulation.The resulting reduction in northward heat transport may have further altered latitudinal temperature and moisture gradients.

Concluding remarks (2.7.6)

Natural climate variability in the Arctic over the past two million years has been large. In particular, the past 20000-year period is now known to have been highly unstable and prone to rapid changes, especially temperature increases that occurred rapidly (within a few decades or less).These temperature increases occurred during glacial terminations and at the onset of D-O interstadials. This instability implies rapid, closely linked changes within the earth’s environmental system, including the hydrosphere, atmosphere, cryosphere, and biosphere. Not only has the climate of the Arctic changed significantly over the past two million years, there have also been pronounced regional variations associated with each change.

The Arctic is not homogeneous and neither is its climate, and past climate changes have not been uniform in their characteristics or their effects. Many of these changes have not been synchronous nor have they had equal magnitudes and rates of change. Climate changes in one part of the Arctic may trigger a delayed response elsewhere, adding to the complexity. The paleoenvironmental evidence for the Arctic suggests that at certain times, critical thresholds have been passed and unpredictable responses have followed. The role that anthropogenic changes to the climate system might play in exceeding such thresholds and the subsequent response remains unclear.

It is clear that between 400 and 100 years BP, the climate in the Arctic was exceptionally cold. There is widespread evidence of glaciers reaching their maximum post-Wisconsinan positions during this period, and the lowest d18O values and melt percentages for at least 1000 years are recorded in ice cores for this interval. The observed warming in the Arctic in the latter half of the 20th century appears to be without precedent since the early Holocene[125].

Chapter 2: Arctic Climate - Past and Present

2.1 Introduction to Arctic climate: Past and Present (Arctic climate variability prior to 100 years BP)
2.2 Arctic atmosphere
2.3 Marine Arctic
2.4 Terrestrial Water Balance in the Arctic
2.5 Influence of the Arctic on global climate
2.6 Arctic climate variability in the twentieth century
2.7 Arctic climate variability prior to 100 years BP
2.8 Summary and key findings of ACIA on Arctic Climate - Past and Present

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Citation

Committee, I. (2012). Arctic climate variability prior to 100 years BP. Retrieved from http://editors.eol.org/eoearth/wiki/Arctic_climate_variability_prior_to_100_years_BP
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