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Paleoclimate History of the ArcticG H Miller, University of Colorado, Boulder, CO, USA; University of Iceland, Reykjavik, IcelandJ Brigham-Grette, University of Massachusetts, Amherst, MA, USAR B Alley, Pennsylvania State University, University Park, PA, USAL Anderson, US Geological Survey, Denver, CO, USAH A Bauch, Mainz Academy of Sciences, Humanities, and Literature, Kiel, GermanyM S V Douglas, University of Alberta, Edmonton, AB, CanadaM E Edwards, University of Southampton, Southampton, UKS A Elias, Royal Holloway, University of London, Egham, UKB P Finney, Idaho State University, Pocatello, ID, USAJ J Fitzpatrick, US Geological Survey, Denver, CO, USAS V Funder, University of Copenhagen, Copenhagen, DenmarkA Geirsdottir, University of Iceland, Reykjavik, IcelandT D Herbert, Brown University, Providence, RI, USAL D Hinzman, University of Alaska Fairbanks, Fairbanks, AK, USAD S Kaufman, Northern Arizona University, Flagstaff, AZ, USAG M MacDonald, University of California, Los Angeles, CA, USAL Polyak, The Ohio State University, Columbus, OH, USAA Robock, Rutgers University, New Brunswich, NJ, USAM C Serreze, University of Colorado, Boulder, CO, USAJ P Smol, Queen’s University, Kingston, ON, CanadaR Spielhagen, Leibniz Institute for Marine Sciences, Kiel, GermanyJ W C White, University of Colorado, Boulder, CO, USAA P Wolfe, University of Alberta, Edmonton, AB, CanadaE W Wolff, British Antarctic Survey, Cambridge, UK

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Early Cenozoic Warm Times

At the start of the Cenozoic, the planet was ice-free; there wasno sea ice in the Arctic Ocean and there was no Greenland orAntarctic ice sheet. Benthic Foraminifera d18O document along-term cooling of the global deep ocean during the Ceno-zoic (Zachos et al., 2001). Because oceanic bottom watersoriginate in the polar oceans, Arctic climate history is broadlyconsistent with the global deep-ocean record. Although generalcooling and an increase in terrestrial ice volume characterizethe Cenozoic, this first-order was punctuated by short- andlonger-lived climate reversals, by variations in cooling rate,and by additional features related to growth and shrinkage ofice once ice sheets were well established. A detailed ArcticOcean record that is equivalent to the global record is not yetavailable. Because the Arctic Ocean is connected to the globalocean only by a single narrow deep channel and some inter-mittently exposed shallow shelves, the Arctic Ocean d18O re-cord often differs from that of the global ocean, hinderingcorrelations. Emerging paleoclimate reconstructions fromthe Arctic Ocean derived from sediments recovered from theLomonosov Ridge (Polyak et al., 2010) shed new light on theCenozoic evolution of the Arctic Basin, but the data have yet tobe fully integrated with evidence from high-latitude terrestrialrecords or with the other records from elsewhere in the ArcticOcean. The presence of trace amounts of coarse clastic debrismay indicate early onset of Arctic Ocean sea ice or calving landice, but the evidence remains circumstantial and alternativeexplanations are possible.

General cooling through the Cenozoic is attributed mainlyto a slow drawdown of greenhouse gases in the atmospherethrough the weathering of silicic rocks that exceeded the releaseof stored carbon through volcanism and reprocessing. Manylines of proxy evidence show that atmospheric CO2 was higherin the Cretaceous than it has been recently, and that it subse-quently fell in parallel with the cooling. Changes in continen-tal positions, sea level, and oceanic circulation as well asbiological evolution also may have contributed, but are lesserfactors.

Pliocene Warmth and the Transition into QuaternaryIce-Age Cycles

During the middle Pliocene (!3.5 Ma), forests occupied largeregions near the Arctic Ocean that are currently polar deserts.Fossils of the marine bivalve Arctica islandica, which does notlive where there is seasonal sea ice, found in marine deposits asyoung as 3.2 Ma old on Meighen Island at 80"N, likely recordthe peak Pliocene warmth of the ocean (Fyles et al., 1991).Multiple temperature proxies from terrestrial peat in the highCanadian Arctic indicate mean annual Pliocene temperatures19 "C higher than present (Ballantyne et al., 2010); other siteshave produced large temperature differences for the Pliocene ofthe North American Arctic, although slightly lower magnitudes(14 "C, Ellesmere Is., Ballantyne et al., 2006; beetles and plantsindicative of summers as much as 10 "C warmer than recentlymeasured temperatures and even larger wintertime warming to

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Encyclopedia of Quaternary Science, (2013), vol. 3, pp. 113-125

a maximum of 15 "C or more, Elias and Matthews, 2002).Although continental configurations were similar to modernforests extended to the Arctic shoreline, nearly eliminating theArctic tundra biome, and sea level stood !25 m higher thanpresent, implying much smaller Greenland (possibly limited tosmall mountain ice caps and glaciers) and Antarctic ice sheets(possibly without a West Antarctic Ice Sheet).

Explaining the warmth of the Arctic during the Plioceneremains enigmatic. The duration of warmth exceeds cycle-lengths related to orbital irregularities (>105 years), and conti-nental configurations and gateways were similar to the presentday. It is likely that the ocean transported more heat intothe Arctic from the Atlantic during peak Pliocene warmth(Haywood et al., 2009), although the status of Arctic Oceansea ice remains uncertain, as does the amount of ice on Green-land. Best estimates of atmospheric CO2 loadings are similar topresent (AD 2011) levels (Pagani et al., 2010), although mostclimate models driven by changes in greenhouse gases (mostlyCO2) tend to underestimate Arctic warmth (e.g., Sloan andBarron, 1992). A consensus view is that greater atmosphericand/or oceanic meridional heat transport, coupled with somecombination of stronger greenhouse gas forcing and strongerlong-term positive feedbacks than those found in many climatemodels, are required to explain the reconstructed temperaturesfor the Arctic during the warmest intervals of the Pliocene.

The Early Quaternary

The development of large continental ice sheets in the NorthAmerican and Eurasian Arctic marks the onset of the Quater-nary. The first secure evidence for the growth of these ice sheetsis the appearance of ice-rafted debris (IRD) in the central NorthAtlantic between 3.0 and 2.5 Ma, far from shores where sea-sonal sea ice may transport sediment. Recently, the Interna-tional Union of Geological Sciences has formalized a revisionof the geological timescale that places the base of the Quater-nary Period at 2.58 Ma, coincident with the Gauss–Matuyamapaleomagnetic reversal (Gibbard et al., 2010), and closelyaligned to the start of the ice-age cycles. The full Quaternaryis best represented in deep-sea sediment, where the changes ind18O of benthic Foraminifera are used to subdivide the Qua-ternary into 104 numbered marine isotopic stages (MIS),counting consecutively back from the present (Lisiecki andRaymo, 2005), with odd-numbered stages representing globalwarm times/minimum ice volumes, and even-numbered stagesrepresenting peak ice volumes.

The appearance of large, quasistable continental ice sheetscoincides approximately with the start of the Quaternary, butgeneral cooling through the Cenozoic is reflected by the ap-pearance of smaller glaciers and ice caps much earlier, duringthe Pliocene and possibly earlier. Although the pattern ofglacial onset over Greenland remains uncertain, small amountsof IRD are found in marine sediment recovered from thecontinental shelf around Greenland as old as 7 Ma (St. Johnand Krissek, 2002). IRD in proximal marine sediment, but notin the central Atlantic, likely indicates that climate cooledsufficiently during the Pliocene and possibly even as early asthe late Miocene, to allow small ice caps and glaciers to form athigh elevations on Greenland. Outlet glaciers delivering small

icebergs to the sea during the coldest intervals likely explainsthe observed IRD, but the total ice volume on Greenland mayhave been very small. Indeed, the first significant appearanceof likely Greenland-derived IRD in the open ocean is observedat !3.2 Ma (Kleiven et al., 2002), which may represent theinception of a Greenland Ice Sheet, although the geographicdistribution suggests Scandinavia and North America are alsopotential sources. Widespread IRD in the North Atlantic beginsabout 2.7 Ma (Kleiven et al., 2002), at about the onset of theQuaternary.

The best terrestrial record of the early phases of late Ceno-zoic ice growth comes from Iceland, where unusually highvolcanic productivity, caused by the superposition of mantle-plume and mid-ocean ridge magmatism, resulted in anenvironment where constructive geologic processes exceeddestructive processes. This unusual state results in the uniquepreservation of glacial deposits in Iceland extending well backinto the Pliocene. Because lava flows can only occur if thelandscape is ice-free, the general consensus is that such flowsrepresent interglacials, interstadials, or other ice-free periods,whereas hyaloclastites (subglacially formed volcanics) and tillrepresent glaciations. Consequently, a record of past glacia-tions (often including ice-flow features) and interglaciationsis preserved within stacked volcanic sequences. Mapping anddating of remnants of glacial deposits found imbeddedwithin the lava flows of Iceland have revealed at least 22glacial–interglacial cycles during the last !3 Ma (Geirsdottiret al., 2011). Four phases of glaciation have been identified:(1) The inception phase, between >4 and 3 Ma, when ice capswere concentrated over the mountainous regions in the south-east of Iceland. (2) A transitional growth phase (3.0–2.5 Ma),when the ice sheet(s) extended their margins to the north andwest, reaching the sea in the north by Tjornes by 2.5 Ma.(3) The presence of glacial deposits across all of the highlandsby 2.5–2.4 Ma, suggests that progressive cooling resulted incontinued incremental north and westward growth of theQuaternary ice sheet from its nucleus in southeast Iceland.By this time, the ice sheet covered more than half of the island.(4) Shortly afterwards, between 2.5 and 2.2 Ma, a single icesheet covered the whole of Iceland for the first time, reachingthe modern shoreline, and full-scale glacial–interglacial cyclic-ity was established (Geirsdottir, 2011).

For the first 1.5–2.0 Ma of the Quaternary, cycles of waxingand waning continental glaciation appeared at 41 ka intervals,as the climate oscillated between glacial and interglacial states.Among the prominent but interglaciations during this period isone at about 2.4 Ma that is especially well represented by theKap København Formation, a 100-m-thick sequence of estua-rine sediments that covered an extensive lowland area near thenorthern tip of Greenland (Funder et al., 2001). The rich andwell-preserved fossil fauna and flora in the Kap KøbenhavnFormation record warming from cold conditions as intergla-ciation, followed by cooling during the subsequent 10–20 ka.During the peak warmth, forest trees reached the Arctic Oceancoast, 1000 km north of the northernmost trees today. Basedon this warmth, Funder et al. (2001) suggested that theGreenland Ice Sheet must have been reduced to local ice capsin mountain areas. Although finely resolved time records arenot available throughout the Arctic Ocean of that time,by analogy with present faunas along the Russian coast,

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the coastal zone would have been ice-free for 2–3 months insummer. Today, this coast of Greenland experiences year-round sea ice, and models of diminishing sea ice in a warmingworld generally indicate long-term persistence of summertimesea ice off these shores (e.g., Holland et al., 2006). Thus, thereduced sea ice off northern Greenland during deposition ofthe Kap København Formation suggests a widespread intergla-ciation during which Arctic sea ice was greatly diminished, andpossibly absent in summer.

During the warm interval described above, precipitationwas greater and temperatures were higher than at the peak ofthe current interglaciation, about 8 ka, and the temperaturedifference was larger during winter than during summer.Higher temperatures were not caused by notably greatersolar insolation, owing to the relative repeatability of theMilankovitch variations over millions of years (e.g., Bergeret al., 1992). Uncertainties in atmospheric CO2 concentra-tion, ocean heat transport, and perhaps other factors at thetime of the Kap København Formation are sufficiently large topreclude strong conclusions about the causes of the unusualwarmth.

Potentially correlative records of warm Early Quaternaryinterglacial conditions are found in deposits along coastalplains in northern and western Alaska. Interglacial marinetransgressions that repeatedly flooded the Bering Strait mod-ified the configuration of adjacent coastlines, altered regionalcontinentality, and reinvigorated the exchange of watermasses between the North Pacific, Arctic, and North Atlanticoceans. Since the first submergence of the Bering Strait about5.5–5.0 Ma (Marincovich and Gladenkov, 2001), this marinegateway allowed relatively warm Pacific water to reach as farnorth as the Beaufort Sea (Brigham-Grette and Carter, 1992).The Gubik Formation of northern Alaska records at least threewarm high sea stands in the early Quaternary. During theColvillian transgression, about 2.7 Ma, the Alaskan CoastalPlain supported open boreal forest or spruce–birch woodlandwith scattered pine and rare fir and hemlock (Nelson andCarter, 1991). Warm marine conditions are confirmed bythe general character of the ostracode fauna, which includesPterygocythereis vannieuwenhuisei (Brouwers, 1987), an extinctspecies of a genus whose modern northern limit is theNorwegian Sea and which, in the northwestern AtlanticOcean, is not found north of the southern cold-temperatezone (Brouwers, 1987). Despite the high sea level and relativewarmth indicated by the Colvillian transgression, erraticboulder in Colvillian deposits southwest of Barrow, Alaska,indicate that glaciers then terminated in the Arctic Ocean andproduced icebergs large enough to reach northwest Alaska.

The subsequent Bigbendian transgression (about 2.5 Ma),characterized by rich molluscan faunas including the gastro-pod Littorina squalida and the bivalve Clinocardium californiense(Carter et al., 1986), indicates renewed warmth along northernAlaska. The modern northern limit of both of these molluskspecies is well to the south (Norton Sound, Alaska). The pres-ence of sea otter bones suggests that the limit of seasonal ice onthe Beaufort Sea was restricted during the Bigbendian intervalto positions north of the Colville river and thus well north oftypical twentieth century positions (Carter et al., 1986); mod-ern sea otters cannot tolerate severe seasonal sea-ice conditions(Schneider and Faro, 1975).

The Fishcreekian transgression (about 2.4–2.1 Ma) hasbeen correlated with the Kap København Formation onGreenland (Brigham-Grette and Carter, 1992). However, agecontrol is imprecise, and Brigham (1985) and Goodfriend et al.(1996) suggested that the Fishcreekian could be as young as1.4 Ma. This deposit contains several mollusk species that cur-rently are found only south of the winter sea ice margin. More-over, sea otter remains and the intertidal gastropod L. squalida atFish Creek suggest that perennial sea ice was absent or severelyrestricted during the Fishcreekian transgression (Carter et al.,1986). Correlative deposits rich inmollusk species that currentlylive only well to the south are reported from the coastal plain atNome, Alaska (Kaufman and Brigham-Grette, 1993).

The available data clearly indicate episodes of relativelywarm conditions that correlate with high sea levels andreduced sea ice in the early Quaternary. The high sea levelssuggest the melting of land ice. Thus, the correlation of warmthwith diminished ice on land and at sea is analogous to feat-ures of contemporary change, as indicated by recent instru-mental observations, model results, and data from other timeintervals.

The Mid-Pleistocene Transition: 41 and 100 ka Worlds

Throughout the Quaternary the cyclical waxing and waning ofcontinental ice sheets has dominated global climate variabilityin response to variations in solar insolation caused by featuresof Earth’s orbit. After the onset of glaciation in North America,ice growth and decay occurred on the same frequency as theEarth’s obliquity (tilt), a 41 ka cycle. But between 1.2 and0.7 Ma, the variations in ice volume became larger and theperiods much longer, averaging approximately 100 ka betweeninterglaciations by 700 ka. Although Earth’s eccentricity varieswith an approximately 100 ka period, this variation results inonly small changes in solar insolation in the key regions of icegrowth, far less than precession and obliquity cycles, so eccen-tricity in unlikely to be the cause of this change; the reason forthe dominant 100 ka period in ice volume remains debated.

The MPT is of particular interest because it was not causedby any major change in Earth’s orbital behavior, and so thetransition likely reflects a fundamental threshold within theclimate system. Models for the 100 ka variability commonlyassign a major role to the ice sheets themselves and especiallyto the Laurentide Ice Sheet, which dominated the total globalchange in ice volume. Roe and Allen (1999) assessed six dif-ferent explanations of this behavior and found that all fit thedata rather well; no one model has superiority over another.

A key observation that must be satisfied in any explanationfor the mid-Pleistocene transition (MPT) is that the ice sheetsof the last 700 ka were larger in volume but smaller in areathan were the largest early Quaternary ice sheets (Balco et al.,2005a,b). Clark and Pollard (1998) used this observation toargue that the early Laurentide Ice Sheets must have had asubstantially lower surface elevation than in the late Pleisto-cene, possibly by as much as 1 km. They argued that at theonset of the Quaternary, ice sheets advanced over easily de-formed water-saturated weathered terrain. The low basal shearstrength allowed fast ice flow, producing aerial extensive, butrelatively thin ice sheets. Successive glacial cycles gradually

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eroded the regolith, eventually exposing irregular bedrock inthe central region. With increased resistance to basal sliding,the ice sheets eventually became thicker with a steeper surfaceprofile, but not quite as extensive aerially as the early Quater-nary ice sheets, despite their greater ice volume.

Other hypotheses emphasize the gradual global coolingthat began in the early Cenozoic and continued throughmuch of the Quaternary (Raymo et al., 2006; Ruddiman,2003). For example, cooling may have passed a thresholdwhereby ice sheets attained sufficient thickness in theirgrowth phase that the duration of the ‘warm’ portion of the41 ka tilt cycle was too short to melt enough ice to causedeglaciation. Then the 100-ka cycle may result from the cu-mulative build up of the Laurentide Ice Sheet until it wassufficiently thick that it finally trapped enough of Earth’sinternal heat to thaw the ice-sheet bed, and the consequentfaster ice flow lead to deglaciation. Marshall and Clark (2002)modeled the growth and decay of the Laurentide Ice Sheetand found that during growth the ice was frozen to its bed,decreasing flow velocities. After many tens of thousands ofyears, the modeled ice had thickened sufficiently to trapenough geothermal heat to thaw the bed, which allowedfaster flow. Faster flow of the ice sheet lowered the surfacegradient, which allowed warming and melting. Behavior suchas that described could cause the main variations of ice vol-ume to be slower than the orbitally driven variations in inso-lation. Alternative hypotheses require interactions in theSouthern Ocean between the ocean and sea ice and betweenthe ocean and the atmosphere (Gildor et al., 2002). Forexample, Toggweiler (2008) suggested that because of theclose connection between the southern westerly winds andmeridional overturning circulation in the Southern Ocean,shifts in wind fields could control the exchange of CO2 be-tween the ocean and the atmosphere. Carbon models supportthe notion that weathering and the burial of carbonate can beperturbed in ways that alter deep ocean carbon storage andresult in 100 ka CO2 cycles. Others have suggested that 100 kacycles and CO2 might be controlled by variability in obliquitycycles (i.e., two or three 41 ka cycles that produce a mixture of80 and 120 ka cycles after 700 ka; Huybers, 2006, 2007, or byvariable precession cycles altering the 19 and 23 ka cycles;Raymo, 1997). Ruddiman (2006) furthered these ideas butsuggested that since 900 ka, CO2-amplified ice growth con-tinued at the 41 ka intervals but that polar cooling dampe-ned ice ablation. His CO2-feedback hypothesis suggests amechanism that combines the control of 100 ka cycles withprecession cycles (19 and 23 ka) and with tilt cycles (41 ka).

Other hypotheses also exist for these changes. A completeexplanation of the onset of extensive glaciation on NorthAmerica and Eurasia at the beginning of the Quaternary, andof the MPT from 41 to 100 ka ice-age cycles, remain the objectof ongoing investigations.

A Link Between Ice Volume, Air Temperature, andGreenhouse Gases

The globally averaged temperature decrease during the 100-kaice-age cycles since the MPT was about 5–6 "C (Jansen et al.,2007). Larger decreases are calculated for the Arctic and close

to the ice sheets, with a decrease of !22 "C atop the GreenlandIce Sheet at the Last Glacial Maximum (LGM; Cuffey et al.,1995). The total change in solar radiation reaching the planetduring these cycles was near zero; orbital features servedprimarily to change insolation seasonally and geographically.

Many factors probably contributed to this large temperaturechange despite the small global change in total insolation.Cooling produced growth of reflective ice that increased theplanetary albedo, while the great height of the ice sheets on itsown reduced Arctic temperatures as well. But changes in atmo-spheric greenhouse gases are necessary to explain the magni-tude of the global cooling.

Antarctic ice cores now provide us with accurate estimatesof atmospheric trace gas concentrations for the past 800 ka(Luthi et al., 2008, and references therein). Complex changes,especially in the ocean, reduced atmospheric carbon dioxideconcentrations during glaciations, and both oceanic and ter-restrial changes reduced atmospheric water vapor, methane,and nitrous oxide, all of which are greenhouse gases.The changes in water vapor and carbon dioxide were mostimportant in the planetary energy balance. Increasing aridityin some areas and glacial outwash in others produced addi-tional dust that reduced the flux of insolation reaching theplanet’s surface (e.g., Mahowald et al., 2006). Cooling causedregions formerly forested to give way to grasslands, tundra ordune fields, further increasing the planetary albedo. WhileEarth’s orbital features paced the ice-age cycles, strong positivefeedbacks are required to provide quantitatively accurate ex-planations of the observed glacial and interglacial changes.

The relation between climate and carbon dioxide has beenstrongly correlated for at least the past 800 ka (Luthi et al.,2008), and the growth and shrinkage of ice, cooling andwarming of the globe, and other changes have repeated alongsimilar, although not identical paths.

Marine Isotopic Stage 11 – A Long Interglaciation

Following the MPT, the growth and decay of ice sheets fol-lowed a 100 ka cycle: brief, warm interglaciations lasted 10 toabout 40 ka, after which ice progressively increased in volumeto a maximum, before ice volumes decreased rapidly and theplanet transitioned into the next interglaciation. Although this100 ka cycle is unlikely to be linked to changes in the 100 kaeccentricity cycle because it produces so little a change in solarenergy reaching the Earth, there is an additional !400 ka or-bital cycle; Earth’s orbit goes from almost round to moreeccentric to almost round in about 100 ka, but the maximumeccentricity reached in these 100-ka cycles increases and de-creases within a 400-ka cycle (Berger and Loutre, 1991). Whenthe orbit is almost round, there is little effect from Earth’sprecession. But, about 400 ka, during MIS 11, the 400-kacycle caused a nearly round orbit to persist, and the interglacia-tion as recorded in both ice cores and marine sediment lastedlonger than previous or subsequent warm times (Jouzel et al.,2007). When Earth’s orbit is nearly round, there is little changein insolation during a precession cycle. Consequently, summerinsolation at high northern latitudes may not have becomelow enough at the end of the first 11 ka precession hemicycleto allow ice growth in the Arctic.




Mid- and low-latitude paleoclimate records show that MIS11 lasted !30 ka, rather than the typical 10 ka duration, butsediments containing MIS 11 are rare in the Arctic and theirtemperature signals are inconclusive (see Helmke and Bauch,2003, and references therein). Sea level during MIS 11 washigher than at any time since (cf. Olson and Hearty, 2009),and data from Greenland are consistent with notable shrinkageor loss of the ice sheet accompanying the longwarmth, althoughthe dating is poorly constrained (Willerslev et al., 2007).

Marine Isotopic Stage 5e: The Last Interglaciation

The warmest millennia of at least the past 250000 years oc-curred during MIS 5, and especially during the warmest partof that interglaciation, MIS 5e (Bauch et al., 1999; Fronvaland Jansen, 1997; Kukla, 2000; McManus et al., 1994). Dur-ing MIS 5e Earth’s orbital parameters aligned to produce astrong positive anomaly in solar radiation during summerthroughout the Northern Hemisphere (Berger and Loutre,1991). Between 130 and 127 ka, the average solar radiationduring the key summer months (May, June, and July) wasabout 11% greater than solar radiation at present throughoutthe Northern Hemisphere, and a slightly greater anomaly,13%, existed over the Arctic. Greater solar energy in summerlead to melting of the Laurentide and Fennoscandinavian icesheets earlier, during the precession cycle than during the lastdeglaciation. Sea level was higher than present 129 ka, coin-cident with the Northern Hemisphere insolation peak,whereas sea level only reached present in the Holoceneabout 6 ka, 5 ka after the peak in Northern Hemisphere inso-lation (CAPE Last Interglacial Project Members, 2006). Thelarge insolation anomaly and intensification of the NorthAtlantic Drift combined to reduce Arctic Ocean sea ice,allow expansion of boreal forest to the Arctic Oceanshore throughout large regions, reduce permafrost, and meltalmost all glaciers in the Northern Hemisphere (CAPE LastInterglacial Project Members, 2006).

High solar radiation in summer during MIS 5e, amplifiedby key boundary-condition feedbacks (especially sea ice, sea-sonal snow cover, and atmospheric water vapor), collectivelyproduced summer temperature anomalies 4–5 "C above presentover most Arctic lands (CAPE Last Interglacial Project Members,2006), substantially above the average Northern Hemispheresummer temperature anomaly (1#1 "C above present). MIS5e demonstrates the strength of positive feedbacks on Arcticwarming (Miller et al., 2010b).

Terrestrial MIS 5e Records

Summers throughout the Arctic were warmer during MIS 5ethan at present, but the magnitude of warming differed spa-tially. Positive summer temperature anomalies were largestaround the Atlantic sector, where summer warming was typi-cally 4–6 "C higher than present. This anomaly extended intoSiberia, but it decreased from Siberia westward to the Europeansector (0–2 "C), and eastward toward Beringia (2–4 "C). TheArctic coast of Alaska had sea-surface temperatures 3 "C aboverecent values and considerably less summer sea ice thanrecently, but much of interior Alaska had smaller anomalies

(0–2 "C) that probably extended into western Canada.In contrast, northeastern Canada and parts of Greenland hadsummer temperature anomalies of about 5 "C, and perhapsmore. A stratified lacustrine sequence from the easternCanadian Arctic that captures MIS 7, MIS 5e, and the Holo-cene, suggests that for at least this portion of the Arctic, onlyfull interglacials were warm enough for lakes to be ice-free insummers long enough for the preservation of biotic-bearingsediment; the record also suggested that twentieth centurywarming represents a no-analog situation for the lake (Axfordet al., 2009).

Precipitation and winter temperatures are more difficultto reconstruct for MIS 5e than are summer temperatures.In northeastern Europe, the latter part of MIS 5e was charac-terized by a marked increase in winter temperatures. A largepositive winter temperature anomaly also occurred in Russiaand western Siberia, although the timing is not as well con-strained (Funder et al., 2002; Gudina et al., 1983; Troitsky,1964). Qualitative precipitation estimates for most other sec-tors indicate wetter conditions than in the Holocene.

Marine MIS 5e Records

Low sedimentation rates in the central Arctic Ocean and therare preservation of carbonate fossils limit the number of sitesat which MIS 5e can be reliably identified in sediment cores.Peak concentrations of a foraminifer species that usually dwellsin subpolar waters were found in MIS 5e zones and interpretedto indicate relatively warm interglacial conditions and muchreduced sea-ice cover in the interior Arctic Ocean (Adler et al.,2009; Nørgaard-Pedersen et al., 2007a,b). Interpretation ofthese and other proxies is complicated by the strong verticalstratification in the Arctic Ocean; today, warm Atlantic water(>1"C) is in most areas isolated from the atmosphere by arelatively thin layer of cold (<$1 "C) low-salinity surfacewater, limiting the transfer of heat to the atmosphere. It isnot always possible to determine whether subpolar Foraminif-era found in Arctic Ocean sediment cores lived in warm watersthat remained isolated from the atmosphere below the coldsurface layer, or whether the warm Atlantic water had displacedthe cold surface layer and was interacting with the atmosphereand affecting its energy balance.

Landforms and fossils from the western Arctic and BeringStrait indicate substantially reduced sea ice during MIS 5. Thewinter sea-ice limit is estimated to have been as much as800 km farther north than its average twentieth century posi-tion, and summer sea ice was likely to have been muchreduced relative to present (Brigham-Grette and Hopkins,1995). These reconstructions are consistent with the north-ward migration of treeline by hundreds of kilometersthroughout much of Alaska and the expansion of treeline tothe Arctic Ocean coast in the Far East of Russia (Lozhkin andAnderson, 1995).

Marine Isotopic Stage 3 (70–30 ka): Rapid ClimateOscillations

The Arctic climate history through MIS 3 is difficult to recon-struct because of the paucity of continuous records and the




difficulty in providing a secure time frame. Continental icesheets were reduced, relative to ice volumes during MIS 2 andMIS 4, and sea level was slightly higher, but the coastline wasstill well offshore in most places. The d18O record of tempera-ture over the Greenland Ice Sheet and other ice-core data showthat the North Atlantic region experienced repeated abruptclimate oscillations between fully glacial and interstadial con-ditions throughout MIS 3. These oscillations, known as Dans-gaard–Oeschger (DO) events, each lasted several hundred to afew thousand years, with temperatures changing by as much as15 "C (Alley, 2007, and references therein). Similar climateexcursions are also recorded in cave sediments from Chinaand Yemen (Burns et al, 2003; Dykoski, et al., 2005; Wanget al., 2001), in high-resolution marine records off California(Behl and Kennett, 1996), and sediment cores from the Car-iaco Basin (Hughen et al., 1996), the Arabian Sea (Schulz et al.,1998) and the Sea of Okhotsk (Nurnberg and Tiedemann,2004), among many other sites. Rapid climate changes withinMIS 3 were thus hemispheric to global in nature and areconsidered a sign of large-scale coupling between the cryso-phere, ocean, and atmosphere. Although the cause of theseevents is still debated, they were likely the result of abruptreorganizations of the ocean’s thermohaline circulation, prob-ably related to inherent ice sheet instabilities that introducedlarge quantities of icebergs and fresh water into the NorthAtlantic (MacAyeal, 1993).

The terrestrial Arctic vegetation record reflects increasedcontinentality that very likely contributed to relatively highsummer temperatures that presumably were offset by colderwinters. Variable paleoenvironmental conditions were typicalof the MIS 3 (Karaginskii) period throughout Arctic Russia,although regional correlations are difficult. However, evi-dence for repeated abrupt temperature oscillations is weak,possibly because of the difficulty of obtaining precise dates inthis time range. On the New Siberian Islands, Andreev et al.(2001) document the existence of graminoid-rich tundra cov-ering wide areas of the emergent shelf during MIS 3, withsummer temperatures possibly 2 "C higher than during thetwentieth century. At Elikchan 4 Lake in the upper Kolymadrainage, the sediment record contains at least three MIS 3intervals when summer temperatures and treeline reachedlate Holocene conditions (Anderson and Lozhkin, 2001).MIS 3 insect faunas in the lower Kolyma are thought to havethrived in summers that were 1–4.5 "C higher than recently(Alfimov et al., 2003).

The warmest widespreadMIS 3 interval in Beringia occurred44–35 ka (Anderson and Lozhkin, 2001); it is well representedin proxies from interior sites, although there is little or novegetation response in areas closest to Bering Strait. Summerwarmth appears to have been strongest in eastern Beringiawhere temperatures were only 1–2 "C lower than at presentbetween 45 and 33 ka (Elias, 2007). The warmest interval ininterior Alaska (Fox Thermal Event), about 40–35 ka, wasmarked by spruce forest tundra, although in the Yukon, forestswere most dense a little earlier, about 43–39 ka (Anderson andLozhkin, 2001).

The transition from MIS 3 to MIS 2 is marked by a shiftfrom warm-moist to cold–dry conditions in western Beringia,but is absent or subtle in all but a few records in Alaska(Anderson and Lozhkin, 2001).

Marine Isotopic Stage 2 (30–15 ka): The LGM

The LGM was cold globally, but particularly cold in the Arctic.Planetary temperatures were 5–6 "C lower than at present(Jansen et al., 2007), whereas mean annual temperatures incentral Greenland were depressed more than 20 "C (Cuffeyet al., 1995), with a similar summer temperature reductionestimated for Beringia (Elias et al., 1996). Much of the Arcticlands lay beneath continental ice sheets, and the Arctic Oceanwas mantled by a continuous cover of sea ice and entrappedicebergs (Bradley and England, 2008). The lack of ocean–atmosphere transfer of heat during winter months would resultin exceptionally cold winters, not only over the Arctic Oceanproper, but also across adjacent lands. These factors togetherlimit the direct records of LGM climate across most of the Arctic.The available evidence for terrestrial ecosystems suggests thatlow- and high-shrub tundra were extremely restricted during theLGM, and largely replaced by polar desert or graminoid and forbtundra and an intergradation of steppe and tundra in the northEurasian interior (Bigelow et al., 2003), ecosystem changesconsistent with cold, dry climates with little snow cover.

Global ice volumes peaked about 21 ka. Rising solar inso-lation across the Northern Hemisphere in summer and in-creases in greenhouse gases caused Northern Hemisphere icesheets to begin to recede from their maxima shortly after 20 ka,with the rate of recession increasing noticeably after about16 ka (see, e.g., Clark et al., 2004; Dyke et al., 2003). Mostcoastlines became ice-free before 12 ka, and ice continued tomelt rapidly as Northern Hemisphere summer insolationreached a peak (!9% above modern) about 11 ka.

A major climate reversal during the deglacial cycle, theYounger Dryas, was initially recognized in pollen recordsfrom northwest Europe, and is now formally defined by thelarge isotopic excursion in the Greenland Ice Sheet cores, be-tween 12.9 and 11.7 ka (Rasmussen et al., 2006). It can beconsidered that last of the large DO events, with its strongestexpression centered on the northern North Atlantic, and down-stream into Eurasia, but also effecting to a lesser degree the restof the Northern Hemisphere.

Marine Isotopic Stage 1, The Holocene

The transition from MIS 2 to MIS 1, which marks the start ofthe Holocene, is commonly placed at the abrupt terminationof the Younger Dryas cold event, about 11.7 ka (Rasmussenet al., 2006). By 8 ka the main unstable Northern Hemisphereice sheets had been reduced in size that there impacts on Arcticclimate was minimal, and by 6 ka, sea level and ice volumeswere close to present, and greenhouse gases in the atmospherediffered little from preindustrial conditions (e.g., Jansen et al.,2007). The Holocene is identified as a separate epoch, notbecause the climate of the present interglaciation is particularlynoteworthy, but because it is the period when humans becamesignificant agents of geological change.

Holocene Thermal Maximum

A wide variety of evidence from terrestrial and marine archivesindicates that peak Holocene Arctic summertime warmth was




achieved during the early Holocene, when most regions of theArctic experienced sustained summer temperatures thatexceeded the observed twentieth century values. This periodof peak warmth, is geographically variable in its timing(Kaufman et al., 2004). The ultimate driver of the warmingwas orbital forcing, which produced increased summer solarradiation across the Northern Hemisphere. At 70"N June in-solation 11 ka was about 45 W m$2 larger than present, for atotal change of about 9% (Berger and Loutre, 1991). Januaryinsolation did not change because there is no insolation thatfar north in January. Consequently, the net annual insolationchange for the Arctic during a precession cycle exceeds that oflower latitudes.

High-resolution (decades to centuries) archives containingmultiple climate proxies are available for most of the Holocenethroughout the Arctic. Consequently, the Holocene recordallows evaluation of the range of natural climate variabilityand of the magnitude of climate change in response to rela-tively small changes in forcings. Most Arctic paleoenvironmen-tal records for the Holocene thermal maximum (HTM) recordprimarily summer temperatures. Many different proxies havebeen exploited to derive these reconstructions, including bio-logical proxies (pollen, diatoms, chironomids, dinoflagellatecysts, and other microfossils), elemental and isotopic geo-chemical indices from lacustrine and marine sediments and icecores, borehole temperatures, age distributions of radiocarbon-dated tree stumps north of (or above) current treeline, and theextralimital distributions of thermophilous marine invertebratesand whales.

A synthesis of 140 Arctic paleoclimatic and paleoenviron-mental records extending from Beringia westward to Iceland(Kaufman et al., 2004) outlines the nature of the HTM in thewestern Arctic. The HTM has an average duration of 2.1 ka inBeringia to 3.5 ka in Greenland. The interval from 10 to 4 kacontains the greatest number of sites recording HTM condi-tions and the greatest spatial extent of those conditions in thewestern Arctic. The HTM began earliest in Beringia, wherewarmer-than-present summers became established 14–13 ka.Intermediate ages for HTM initiation (10–8 ka) are apparent inthe Canadian Arctic Archipelago and in central Greenland.The HTM on Iceland occurred a bit later, 8–6 ka, whereas theonset on Svalbard was 10.8 ka. The latest general HTM onset(7–4 ka) was in central Canada. Most regions registered theonset of summer cooling by 6–5 ka. The average summertemperature anomaly during the HTM across the western Arcticwas 1.6 "C, although it ranged from 0.5 to 3.0 "C, with thelargest changes in the North Atlantic sector and the smallestover North Pacific sector.

The spatial pattern of HTM onset can be explained in partby regional variations in the proximity to continental icesheets, which depressed temperatures nearby until the icemelted back, and to changes in freshwater delivery to theNorth Atlantic, which influenced the strength of the Atlanticmeridional circulation.

Sea-ice conditions in the Arctic Ocean and adjacent chan-nels have been developed by from a wide array of key proxiesincluding the remains of whales and walrus, warm-water ma-rine mollusks, changes in microfaunal assemblages, especiallydiatoms and dinoflagellate cysts, and more recently usingbiomarkers diagnostic of sea-ice or sea-ice-free organisms

(Polyak et al., 2010). These reconstructions parallel the terres-trial record for the most part. They demonstrate that an in-creased mass of warm Atlantic water moved into the ArcticOcean at about the beginning of the Holocene. An obligateAtlantic-water mollusk appeared on northern Svalbard by10.6 ka, with peak penetration of Atlantic water between 9.3and 8 ka (Salvigsen, 2002), with most thermophilous mollusksdisappearing from Svalbard by!5 ka. Dates on the flooding ofBering Strait range from !12 ka to slightly before 13 ka, al-though precise dating remains dependent on knowledge of themarine reservoir age at that time. The influx of Pacific water tothe Arctic Ocean peaked between 9 and 5 ka, then diminishedthrough the late Holocene. The extra ocean heat from bothAtlantic and Pacific sources, coupled with increased summerinsolation, decreased the area of perennial sea-ice cover duringthe early Holocene. Decreased sea-ice cover in the westernArctic also may be indicated by changes in concentrations ofsea-salt sodium in the Penny Ice Cap, eastern Canadian Arctic,and the Greenland Ice Sheet.

As summer temperatures increased through the earlyHolocene, North American treeline expanded northward intoregions formerly mantled by tundra, although the northwardextent appears to have been limited to perhaps a few tens ofkilometers beyond its recent position. In contrast, some tree-line species expanded well beyond their current limits in theEurasian Arctic. Dated macrofossils indicate that individuals ofthe tree genera Betula (birch) and Larix (larch) lived far northof the modern treeline across most of northern Eurasia be-tween 11 and 10 ka, and Larix extended beyond its modernlimits as early as 13–12 ka in eastern Siberia (Binney et al.,2009, and references therein). In the Taimyr Peninsula ofSiberia and nearby regions the most northerly limit reachedby trees during the Holocene was more than 200 km north ofthe current treeline. Spruce (Picea) advanced slightly later thanthe other two genera. Interestingly, pine (Pinus), which nowforms the conifer treeline in Fennoscandinavia and the KolaPeninsula of Russia, does not appear to have established ap-preciable forest cover at or beyond the present treeline in theKola Peninsula until around 7 ka. However, quantitative re-constructions of temperature suggest that summer tempera-tures were above modern by 9 ka, and the development ofextensive pine cover at and north of the present treeline ap-pears to have been delayed relative to this warming. Treelinebegan to retreat southward across northern Eurasia 3–4 ka(Binney et al., 2009).

The timing of the HTM in the Eurasian Arctic overlaps thewidest expression of the HTM in the western Arctic, but thetiming of onset and termination in Eurasia shows less variabil-ity than in North America, and the magnitude of the treelineexpansion and retreat is greater in the Eurasian Arctic.

Permafrost was degraded across much of the Russian Arcticduring the HTM. Melting permafrost was widespread in theEuropean north, but across Siberia permafrost partially thawedlocally (reviewed by Astakhov, 1995). Areas south of the ArcticCircle appear to have experienced deep thawing (100–200 mdepth) through the HTM, followed by renewed permafrostexpansion after 3 ka.

Quantitative estimates of HTM summer temperature anom-alies along the northern margins of Eurasia and adjacentislands range from 1 to 3 "C. Sea-surface temperature anomalies




during the HTM were as much as 4–5 "C for the easternNorth Atlantic sector and adjacent Arctic Ocean.

Neoglaciation

Many climate proxies are available to characterize the overallpattern of late Holocene climate change. Although the timingof the HTM around the Arctic is at least partially modulated bythe waning continental ice sheets, the last substantial ice-sheet-dominated climate perturbation was the brief cold excursionabout 8.2 ka. By 8 ka the rapidly receding continental ice sheetshad become small enough that they no longer exerted a stronginfluence on regional climates, and most proxy summer tem-perature records from the Arctic indicate an overall coolingtrend through the late Holocene, as orbital influences steadilyreduced summer insolation across the Northern Hemisphere.

The least ambiguous proxies that reflect the overall decreasein Northern Hemisphere summer temperatures through theHolocene are mountain glaciers and ice caps. Widespread evi-dence for the expansion or regrowth of glaciers after the HTMwas summarized by Porter and Denton (1967), who intro-duced the term Neoglaciation to describe this reversal of thelong deglaciation following the LGM. Although the timingof the onset of Neoglaciation across the Arctic is poorlyconstrained, most regions show evidence of initial glacier ex-pansion or regrowth between 6 and 5 ka. Throughout theHolocene glacier mass balance has been dominantly con-trolled by summer temperature (Koerner, 2005), and theexpansion of glaciers throughout the Arctic is interpreted toreflect a decrease in summer temperatures.

Neoglacial summer cooling is supported by several otherlines of evidence: a reduction in melt layers in the Agassiz IceCap (Koerner and Fisher, 1990), and the decrease in d18Ovalues in ice cores across Greenland and Arctic Canada(Vinther et al., 2009), clear patterns of change in boreholethermometry through the Greenland Ice Sheet; the retreat oflarge marine mammals and warm-water-dependent mollusksfrom the Canadian Arctic; the southwardmigration of the north-ern treeline across central Canada, Eurasia, and Scandinavia(Barnekow and Sandgren, 2001); the expansion of sea-ice coveralong the shores of the Arctic Ocean on Ellesmere Island, inBaffin Bay, and in the Bering Sea; and the shift in vegetationcommunities inferred from plant macrofossils and pollenaround the Arctic, including Wrangel Island (Lozhkin et al.,2001). The assemblage of microfossils and the stable isotoperatios of Foraminifera in marine sediment indicate a shift towardcolder, lower salinity sea-surface conditions about 5 ka along theEast Greenland Shelf (Jennings et al., 2002) and the westernNordic seas (Koc and Jansen, 1994), suggesting an increasedflux of sea ice from the Arctic Ocean. Where quantitativeestimates of temperature change are available, they generallyindicate that summer temperature decreased by 1–2 "C duringthis initial phase of cooling.

Climate of the Past 2 ka

Climate reconstructions of the past 2 ka have gained increasedattention because they provide a background reference for therange of natural variability against which changes that mightbe tied to anthropogenic activities can be compared. For most

of the Arctic, the pattern of overall summer cooling followingan early HTM forms a multimillennial trend that culminatedin the cold Little Ice Age (LIA) of the thirteenth throughnineteenth centuries. Most Arctic glaciers and ice caps reachedtheir maximum dimensions of the past 8 ka during the LIA(Miller et al., 2010a). High-resolution Arctic climate recon-structions for the first millennium AD are relative rare. Thepan-Arctic (to 60"N latitude) compilation of Kaufman et al.(2009) included 23 proxy temperature records from lake sed-iments, tree rings, and glacier ice that extend back at least1000 years are resolved at annual to subdecadal scale. Thecomposite reconstruction is consistent with other evidencefor millennial-scale trend of overall summer cooling sincethe HTM.

Warmth during Medieval times was initially recognizedfrom evidence in Western Europe (Lamb, 1977), but the termis commonly applied to other regions over a wide temporalrange. We restrict our consideration of Medieval warmth to theperiod between AD 950 and 1250. The strongest evidence forMedieval warmth comes from the northern North Atlanticregion, where there is growing evidence for an episode ofrelatively cold summers between !AD 600 and 950, followedby three centuries of relatively warmer summers during Medi-eval times. Evidence for warmth during this interval is based onreductions in glacier and ice cap dimensions, faunal changes inmarine sediment, speleothems, ice cores, borehole tempera-tures, tree rings, and archaeology. Western Greenland (Crowleyand Lowery, 2000), the Greenland Ice Sheet (Dahl-Jensenet al., 1998), Swedish Lapland (Grudd et al., 2002),northern Siberia (Naurzbaev et al., 2002), Arctic Canada(Anderson et al., 2008), and Iceland (Geirsdottir et al., 2009;Larsen et al., 2011) were all relatively warm around AD 1000.During Medieval time, Inuit populations moved out of Alaskainto the eastern Canadian Arctic and hunted whale from skinboats in regions perennially ice-covered in the twentiethcentury (McGhee, 2004).

Medieval warmth is generally ascribed to a reduction insulfate aerosols derived from explosive volcanism (Jansenet al., 2007), with the extra warmth around the North Atlanticand adjacent regions possibly linked to changes in oceaniccirculation as well (Broecker, 2001).

The Arctic-wide reconstruction of Kaufman et al. (2009)indicates that average summer temperatures were similar toor higher than the warmest interval of the middle ages duringmuch of the first half of the first millennium AD. Following thesixth century and prior to the twentieth, the average warmestinterval occurred between AD 940 and 970. Only some sitesregistered peak preindustrial temperatures during this interval,however. The high average temperature in the composite re-cord during Medieval time is dominated by sites in the north-western North Atlantic sector (Greenland and Canadian HighArctic), especially isotope records from glacier ice. Indicationsof Medieval warmth outside the North Atlantic sector includethe retreat of glaciers in southeastern Alaska (Reyes et al., 2006;Wiles et al., 2008), and the wider tree rings in some high-latitude tree-ring records from Asia and North America(D’Arrigo et al., 2006). However, D’Arrigo et al. (2006) em-phasized the uncertainties involved in estimating Medievalwarmth relative to that of the twentieth century, owing inpart to the sparse geographic distribution of proxy data as




well as to the less coherent variability of tree-growth tempera-ture estimates. Hughes and Diaz (1994) argued that theArctic as a whole was not anomalously warm throughoutMedieval time.

Given the importance of understanding climate of the mostrecent past and the richness of the available evidence, intensivescientific effort has resulted in numerous NorthernHemisphereaverage temperature reconstructions for the past millennium(Mann et al., 2008, and references therein). These reconstruc-tions are based on annually resolved proxy records, primarilyfrom tree rings, with additional records from Greenland icecores and a few annually laminated lake sediment records.Decadally resolved proxy temperature records from the Arctic(poleward of 60"N), primarily from lake sediments, have beenexplored for comparison (Kaufman et al., 2009, and referencestherein). In general, the Northern Hemisphere and the Arcticaverage temperature records are broadly similar: they showmodest summer warmth during Medieval times, a variable,but colder climate from about AD 1250 to 1850, followed bytwentieth century warming as shown by both paleoclimateproxies and the instrumental record. Less is known aboutchanges in precipitation, which is spatially and temporallymore variable than temperature.

The trend toward colder summers after !AD 1250 coin-cides with the onset of the LIA, which persisted until !AD1850, although the timing and magnitude of specific cold in-tervals exhibit regional variability. The climate history of the LIAhas been extensively studied in natural and historical archives,and it is well documented in Europe and North America (Grove,1988). Historical evidence from the Arctic is relatively sparse,but it generally agrees with historical records from northwestEurope. Arctic proxy climate records from glacial and nonglacialsources show that the coldest interval of the Holocene occurredbetween !AD 1450 and !AD 1850, and during this intervalmost glaciers reached their Neoglacial maximum. Recent evi-dence from the Canadian Arctic indicates that, following sub-stantial ice recession in Medieval times, glaciers and ice capsbegan to expand abruptly in the second half of the thirteenthcentury, and that ice expansion was further amplified !AD1450, after which ice caps receded to their pre-LIA marginsonly in recent decades (Miller et al., 2011).

The average summer temperature of the Northern Hemi-sphere during the LIA was no more than 1 "C lower than thetwentieth century average (Mann et al., 2008), but was sub-stantially greater in some regions. LIA cooling appears to havebeen stronger in the Atlantic sector of the Arctic than in thePacific (Kaufman et al., 2004), perhaps because explosive vol-canism promoted the development or transport of sea ice intothe North Atlantic, leading to a reduction in the Atlantic me-ridional overturning circulation, which further amplified LIAcooling there (Miller et al., 2011). Although the initiation ofthe LIA and the structure of climate fluctuations during thismulticentennial interval vary around the Arctic, most recordsshow warming beginning in the late nineteenth century(Kaufman et al., 2009; Overpeck et al., 1997). The end of theLIA was apparently more uniform both spatially and tempo-rally than its initiation.

Summer cooling that marks the transition into the LIA wasin part a consequence of the Holocene-long decrease in North-ern Hemisphere summer insolation due to the orbital changes

described earlier, and triggered by increased explosive volca-nism in the late thirteenth and fifteenth centuries. This transi-tion was likely aided by decreased solar irradiance at thesetimes, and amplified by strong positive feedbacks, especiallyexpanded Arctic Ocean sea ice and terrestrial snow cover, pos-sibly enhanced by a weakening of the meridional circulation inthe northern North Atlantic.

Placing Twentieth Century Arctic Warming in aMillennial Perspective

Much scientific effort has been devoted to learning how twen-tieth and early twenty-first century warmth compares withwarmth during earlier times (e.g., Jansen et al., 2007). A data-model comparison indicates that the long-term cooling trendthat dominated in the Arctic during the past nineteen centurieswas linked to decreasing summer insolation from orbital fac-tors, and that this trend was reversed during the twentiethcentury, despite continued reduction of summer insolationacross the Arctic (Kaufman et al., 2009). Owing to the orbitalchanges affecting midsummer insolation (a drop in Juneinsolation of about 1 W m$2 at 75"N and 2 Wm$2 at 90"Nduring the last millennium; Berger and Loutre, 1991), addi-tional forcing is required to explain why summer temperaturesin recent decades have been similar to, and in most regionswarmer than, summer temperatures achieved in Medievaltimes.

Thin, cold ice caps in the eastern Canadian Arctic preserverooted vegetation beneath them that was killed by the expand-ing ice cover. A recent compilation of 94 14C-dated plantsamples recently emerged from beneath receding ice caps onnorthern Baffin Island shows that many of the ice caps thatformed >1700 years ago persisted through Medieval timesbefore melting early in the twenty-first century (Miller et al.,2011).

Records of surface melting from ice caps offer another viewby which twentieth century warmth can be placed in a millen-nial perspective. The longest record comes from the Agassiz IceCap in the Canadian High Arctic, for which the percentage ofsummer melting of each season’s snowfall is reconstructed forthe past 10 ka. The percent of melt follows the general trend ofdecreasing summer insolation from orbital changes, but somebrief departures are substantial. Of particular note is the signif-icant increase in melt percentage during the past century; cur-rent percentages are higher than any other melt intensity sinceat least 4000 years ago (Fisher et al., 2011).

Changes in lake sediments and their biota record help placethe recent century in a longer perspective. Extensive bioticchanges, especially in the post AD 1850 interval, are inter-preted to reflect recent warming above the Medieval warmthon Ellesmere Island and probably in other regions, althoughalternative explanations have been debated and explored(Smol and Douglas, 2007; Smol et al., 2005). D’Arrigo et al.(2006) show tree-ring evidence from a few North Americanand Eurasian sites that imply that summers were cooler in theMedieval Warm Period than in the late twentieth century,although the statistical confidence is weak. Tree-ring and tree-line studies in western Siberia (Esper and Schweingruber,2004) and Alaska (Jacoby and D’Arrigo, 1995) suggest that




warming since 1970 has enhanced tree growth and follows acircumpolar trend. Hantemirov and Shiyatov (2002) presentrecords from the Yamal Peninsula, Russia, well north of theArctic Circle, showing summer temperatures of recent decadesare the most favorable for tree growth within the past 4 ka.

The National Research Council (2006, p. 3) evaluated theavailable published data on globally and hemispherically av-eraged temperatures for the last millennium, and found that“Presently available proxy evidence indicates that temperaturesat many, but not all, individual locations were higher duringthe past 25 years than during any period of comparable lengthsince A.D. 900.” Greater uncertainties for hemispheric orglobal reconstructions were identified in assessing older com-parisons. Climate models are unable to reproduce the observedwarmth of the late twentieth century unless recent increases ingreenhouse gases due to anthropogenic fossil fuel combustionare included, lending credibility to the argument that latetwentieth century warmth is dominantly an artifact of humanactivities.

Conclusions

As the planet cooled from peak warmth in the early Cenozoic,extensive Northern Hemisphere ice sheets developed by2.6 Ma, leading to changes in the circulation of both the atmo-sphere and oceans. From!2.6 to!1.0 Ma, ice sheets came andwent about every 41 ka, in pace with cycles in the tilt of Earth’saxis, but for the past 700 ka, glacial cycles have been longer,lasting !100 ka, separated by brief, warm interglaciations,when sea level and ice volumes were close to present. Thecause of the shift from 41 to 100 ka glacial cycles is stilldebated. During the penultimate interglaciation, !130to !120 ka, solar energy in summer in the Arctic was greaterthan at any time subsequently. As a consequence, Arctic sum-mers were !5 "C warmer than at present, and almost all Arcticglaciers melted completely except for the Greenland Ice Sheet,and even it was reduced in size substantially from its presentextent. With the loss of land ice, sea level was!5 m higher thanpresent, with the extra melt coming from both Greenland andAntarctica, as well as small glaciers and ice caps. The LGMpeaked !21 ka, when mean annual temperatures over partsof the Arctic were as much as 20 "C lower than at present. Icerecession was well underway 16 ka; by 8 ka even the LaurentideIce Sheet was too small to have a strong influence onthe climate system, and most of the Northern Hemisphereice sheets had melted by 6 ka. Insolation reached a summermaximum (9% higher than at present) !11 ka and has beendecreasing since then, primarily in response to the precessionof the equinoxes. The extra energy elevated early Holocenesummer temperatures throughout the Arctic 1–3 "C abovetwentieth century averages, enough to completely melt manysmall glaciers throughout the Arctic, although the GreenlandIce Sheet was only slightly smaller than at present. EarlyHolocene summer sea-ice limits were substantially smallerthan their twentieth century average, and the flow of Atlanticwater into the Arctic Ocean was substantially greater. As sum-mer solar energy decreased in the second half of the Holocene,glaciers became re-established or advanced, sea ice expanded,and the flow of warm Atlantic water into the Arctic Ocean

diminished. Late Holocene cooling reached its nadir duringthe LIA (!AD 1250 to !AD 1850), when sun-blocking aero-sols from volcanic eruptions and perhaps other causes addedto the orbital cooling, allowing most Arctic glaciers to reachtheir maximum Holocene extent. During the warming of thepast century, glaciers have receded throughout the Arctic, ter-restrial ecosystems have advanced northward, and perennialArctic Ocean sea ice has diminished, with the most pro-nounced pan-Arctic changes occurring since AD 1970.

The histories of temperature and precipitation across Arcticlands during the Cenozoic exhibit clearly defined broad trendsdespite the limitations caused by scattered, and incompleterecords. Arctic climate proxies for the most part reflect summertemperature anomalies, although ice-core isotopes reflect in-fluences from changes in mean annual temperatures, moisturesource area, and seasonality of precipitation. Precipitationchanges in the past are more difficult to reconstruct than arepaleotemperatures, in part due to the greater spatial variabilityof precipitation, but also due to the weak correlation betweenmost proxies and precipitation amount. Nevertheless, somebroad patterns clearly emerge.

On all time scales through the Cenozoic, temperaturechanges have been greater in the Arctic than for the NorthernHemisphere as a whole, for both summer and annual meanchanges. This greater sensitivity reflects the powerful positivefeedbacks acting across the Arctic that amplify changes due toexternal forcings. The strongest feedback on short timescales(decadal to millennial) is the expansion and contraction of seaice, with additional positive feedbacks from changing seasonalsnowcover, changing distribution of forest and tundra ecosys-tem boundaries, growth and decay of permafrost, and changesin the strength of the ocean’s meridional overturning circula-tion. On longer timescales (multimillennial), changes in thelocation and rate of deep convection in the North Atlantic, andthe growth and decay of continental ice sheets are particularlyimportant feedbacks, through changes in energy transfer, CO2

and other greenhouse gas budgets, albedo, and ice-sheetheight, and to a lesser extent, a negative feedback resultingfrom isostatic compensation to the increased ice load thatlowers the ice surface. Accompanying changes in sea-surfacetemperatures control the amount of water vapor released to theatmosphere, influencing the magnitude of the greenhouseeffect.

The great temperature reduction across the Arctic duringglaciations increased the pole-to-equator thermal gradient, in-creasing planetary wind speeds, and the Arctic cold was ampli-fied globally by reduced atmospheric water vapor and CO2, thetwo most important greenhouse gases. Reduced atmosphericwater vapor resulted in globally drier and dustier conditions onaverage, and ice sheet instabilities had dramatic impacts on theoceanic meridional overturning circulation, both with globalconsequences.

The primary forcing that explains the first-order climatetrends through the Cenozoic has been the generally steadydecrease in greenhouse gases in the atmosphere (with somenotable perturbations) that produced a first-order temperaturedecline through most of the Cenozoic. Although this temper-ature decline continued through the Quaternary, high fre-quency temperature variations across the Arctic weresuperimposed on the gradual cooling trend, tied to the growth




and decay of the large continental ice sheets. Throughout theQuaternary, changes in the seasonal and geographic distribu-tion of insolation caused by well-known oscillations in Earth’sorbital parameters, primarily axial tilt (obliquity) and preces-sion of the equinoxes, have paced the expansion and retreat ofNorthern Hemisphere ice sheets, with global climate conse-quences. Arctic Quaternary climates have also responded tovariations in greenhouse gas forcings on glacial–interglacialtimescales. During the Holocene, the general decrease in sum-mer temperatures across the Arctic have been driven by preces-sion, with perturbations to the millennial-scale cooling trendresulting from changes in the frequency of sulfur-rich explosivevolcanic eruptions, and to small changes in solar luminosity.The strong warming trend over the past century across theArctic, and since !1970 in particular, stands in stark contrastto the first-order middle and late Holocene cooling trend, andis very likely a result of increased greenhouse gases that are adirect consequence of anthropogenic activities.

Acknowledgments

This paper is an outgrowth of a series of syntheses commis-sioned by the US Global Change Research Program that wasfocused on paleoclimate of the Arctic. Modified versions ofthese reports were published as a series of five papers in Qua-ternary Science Reviews in 2010. This paper is based on one ofthose papers, “Temperature and precipitation history of theArctic” (Miller et al., 2010a), from which it has been con-densed and revised to include new results since the originalsynthesis was compiled. A wide range of national fundingagencies and logistical support units have supported the re-search from which the reported paleoclimate reconstructionswere derived. The following authors each acknowledge at leastpartial support from the United States National Science Foun-dation as follows: GHM, grants ARC 0714074 and ATM-0318479; RBA, grants 0531211 and 0424589; JWCW, grants0806387, 0537593, and 0519512; LP, grants ARC-0612473and ARC-0806999; LDH, grant OPP-0652838; DSK, grantARC-0455043; MCS, grants ARC- 0531040, ARC-0531302.MSVD, JPS, and APW collectively acknowledge financial sup-port through grants from the Canadian Natural Sciences andEngineering Research Council. Critical logistical support hasbeen provided by the Canadian Polar Continental Shelf Pro-ject, the NSF Office of Polar Programs, and the Icelandic Centerfor Research as well as other national bodies.

See also: Paleoclimate Relevance to Global Warming. Paleoclimate:Introduction; Modern Analog Approaches in Paleoclimatology; TheYounger Dryas Climate Event.

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