Glacial Inceptions: Past and Future

Glacial Inceptions: Past and Future†

Lawrence A. Mysak*

Department of Atmospheric and Oceanic SciencesMcGill University, 805 Sherbrooke Street W.

Montreal QC H3A 2K6

[Original manuscript received 28 August 2007; accepted 4 February 2008]

ABSTRACT The realistic simulation of northern hemisphere glacial inceptions, which occurred during theQuaternary period, has challenged scores of climate theoreticians and modellers. After reviewing theMilankovitch theory of glaciation, a number of earlier modelling studies of the last glacial inception (LGI) whichhave employed either high-resolution General Circulation Models (GCMs) or Earth system Models ofIntermediate Complexity (EMICs) are described. The latter class of models has been developed over the past twodecades in order to investigate the many interactions and feedbacks among the geophysical and biospheric com-ponents of the Earth system that take place over long time scales.

Following a description of the McGill Paleoclimate Model (MPM) and other EMICs, some recent McGill sim-ulations of the LGI in response to orbital (Milankovitch) and radiative (atmospheric CO2) forcings are present-ed. Special attention is given to determining the relative roles of the ocean’s thermohaline circulation, freshwaterfluxes into the ocean, orography, cryospheric processes and vegetation dynamics during the inception phase. Inparticular, it is shown that with the vegetation-albedo feedback included in the model, the buildup of ice sheetsover North America is larger than over Eurasia, in agreement with observations.

This paper concludes with a discussion on the (possible) occurrence of the next glacial period. To address thisissue, which has been inspired by recent publications of Berger and Loutre, MPM simulations of the climate forthe next 100 kyr, forced by various prescribed atmospheric CO2 levels, as well as the future insolation changesas calculated by the Berger algorithm, are presented. The influence of a near-term global warming scenario onglacial inception is also examined. If it is assumed that after such a warming scenario the concentration of CO2in the atmosphere returns to pre-industrial levels (in the range of 280–290 ppm), then the MPM predicts that thenext glacial would start at around 50 kyr after present, which is consistent with the results of Berger and Loutre.Finally, recent simulations of future glacial inceptions using the Potsdam EMIC which includes an atmosphere-ocean carbon cycle component are described. From one of these simulations in which 5000 GtC are releasedinto the atmosphere due to human activities, it is concluded that the current interglacial will last for at leastanother half-million years because of the limited ability of the oceans to absorb such a large carbon release tothe atmosphere.

RÉSUMÉ [Traduit par la rédaction] La simulation réaliste des débuts des périodes glaciaires qui se sontproduites durant le Quaternaire dans l’hémisphère Nord a constitué un défi pour nombre de théoriciens et demodélisateurs du climat. Après un examen de la théorie de Milankovitch sur les glaciations, nous décrivons uncertain nombre d’études de modélisation du début de la dernière période glaciaire (DDPG) faites précédemmentet qui ont utilisé des modèles de circulation générale (MCG) à haute résolution ou des modèles de systèmeterrestre de complexité intermédiaire (EMIC). Cette dernière classe de modèles a été mise au point au cours desdeux dernières décennies dans le but d’étudier les nombreuses interactions et rétroactions se produisant entre leséléments géophysiques et biosphériques du système terrestre sur de grandes échelles de temps.

Après une description du modèle paléoclimatique de McGill (MPM) et d’autres EMIC, certaines simulationsrécentes de McGill du DDPG en réponse aux forçages orbital (Milankovitch) et radiatif (CO2 atmosphérique)sont présentées. Nous nous efforçons en particulier de déterminer les rôles relatifs de la circulation thermohalineocéanique, des flux d’eau douce vers l’océan, de l’orographie, des processus cryosphériques et de la dynamiquede la végétation durant la phase initiale. Nous montrons notamment qu’avec la rétroaction végétation-albédoincluse dans le modèle, l’accumulation de couches glaciaires sur l’Amérique du Nord est plus grande que surl’Eurasie, conformément aux observations.

Cet article se termine par une discussion sur la (possible) prochaine période glaciaire. Pour étudier cettequestion, qui s’inscrit dans la foulée de publications récentes de Berger et Loutre, nous présentons dessimulations du MPM du climat des 100 000 prochaines années, qui sont basées sur un forçage par différentsniveaux de CO2 atmosphérique spécifiés de même que sur les variations futures de l’insolation telles que

ATMOSPHERE-OCEAN 46 (3) 2008, 317–341 doi:10.3137/ao.460303Canadian Meteorological and Oceanographic Society

†This paper is dedicated to the memory of Stephen Mysak, 1906-2007.*Author’s e-mail: [email protected]

mailto:[email protected]

1 IntroductionOver the past few million years, the evolution of the climateon orbital (Milankovitch) time scales has exhibited variousquasi-periodic fluctuations, the most prominent of which arethe relatively recent 100-kyr ice age cycles which can be seenin ice-sheet and deep-ocean sediment core records (Fig. 1).The alternation between relatively short, warm interglacialsand long, cold glacials during the past half-million years isattributed to a complex set of processes that involve orbitalforcing and internal interactions and feedbacks in the climatesystem (Ruddiman, 2001). The main goal of this paper is topresent a review of various attempts to simulate the transitionfrom an interglacial to a glacial period, which is commonlyreferred to as a glacial inception. Both past and future glacialinceptions will be considered. However, since the literatureon glacial inceptions is rather vast, the focus will be on (1) thelast glacial inception (LGI) at around 116 kyr BP (before pre-sent; here ‘present’ is defined as 1950), which occurred dur-ing marine isotope stage (MIS) 5d, and (2) the occurrence ofthe next possible glacial inception. Further, as this paper rep-resents a personal perspective on the topic, the simulationspresented here are mainly from the Earth system modellinggroup at McGill University.

Many of the earlier modelling studies on glacial inceptionhave involved the use of atmospheric General CirculationModels (GCMs) run with a fixed seasonal cycle, a low-levelsummer insolation (i.e., incoming solar radiation at the top ofthe atmosphere) in northern boreal latitudes and a fixed radia-tive (atmospheric CO2) forcing extending over a few decadesor centuries (e.g., see Oglesby (1990), de Noblet et al. (1996),and the references therein). These simulations are calledtimeslice runs. However, since the main driver of the ice agecycles and a glacial inception in particular is the slow changein insolation that takes place over tens of thousands of yearsdue to variations in the Earth’s orbital parameters, recentmodelling studies of these climate changes have employedEarth system Models of Intermediate Complexity (EMICs;Claussen et al., 2002), which are forced by temporally andspatially varying fields that are prescribed for these long timeperiods. Such EMIC simulations are called transient runs,wherein the different components of the climate systemevolve and interact on a variety of time scales in response tothe slowly changing forcing fields.

Among the pioneers who developed EMICs to simulateglacial-interglacial cycles are the many investigators of the

Louvain-la-Neuve group in Belgium (e.g., see Gallée et al.,1992; Loutre and Berger, 2000). However, it is interesting tonote that the first use of the phrase ‘intermediate complexity’in describing a coupled global climate model seems to haveoccurred in Stocker et al. (1992).

The EMIC simulations discussed in this review could beconsidered an extension of earlier transient runs of coupledatmosphere-ice sheet models, in which the atmosphere is rep-resented by an energy balance model (e.g., see Tarasov andPeltier (1999) and the references therein). Since these two-component models do not include an interactive ocean, landsurface, vegetation, or atmospheric moisture components (incontrast to most EMICs), they will not be discussed further inthis review.

With EMICs, it is possible to address many questions thatcannot be answered using GCMs. For example, how andwhen did the last glacial start, over what period of time andwhere did the ice sheets grow substantially in the northernhemisphere? What are the roles of the different climate com-ponents in a glacial inception? Given that the last few inter-glacials (namely, those occurring during MIS 5e, 7 and 9)lasted only on the order of ten thousand years (Fig. 1), it isnatural to ask whether the present interglacial, the 10-kyrHolocene (MIS 1), will end in the not-too-distant future. Overthe past few decades, it has been forecasted (Kukla et al.,1972; Broecker, 1998) that the Holocene would indeed termi-nate soon. More recently, it has been proposed that had it notbeen for the slow anthropogenic emission of greenhousegases thousands of years ago (due to early forest clearance forfarming in Eurasia and rice irrigation in Asia), glaciation innortheastern Canada might have started long before theIndustrial Revolution (Ruddiman, 2003a, 2007). However,the present interglacial may be much longer lasting than thepast few warm periods because of (1) a rather different inso-lation pattern over the next 100 kyr due to the imprint of the400-kyr eccentricity cycle in the Earth’s orbit (see Fig. 3), and(2) the impact of relatively recent anthropogenic activitieswhich have resulted in relatively large greenhouse gas con-centrations in the atmosphere. With regard to (1), it has beenargued that the best analogue for the present interglacial isMIS 11, which started around 410 kyr BP and lasted for aboutthirty thousand years (Fig. 1; EPICA Community Members,2004). (However, Ruddiman (2007) has recently suggested

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calculées au moyen de l’algorithme de Berger. Nous examinons aussi l’influence d’un scénario de réchauffementde la planète dans un avenir rapproché sur l’amorce d’une période glaciaire. Si l’on suppose qu’après un telscénario de réchauffement, la concentration en CO2 atmosphérique revient aux niveaux préindustriels (dans’intervalle 280-290 ppm), alors le MPM prévoit que la prochaine glaciation commencerait dans environ 50 000ans, ce qui correspond aux résultats de Berger et Loutre. Finalement, nous décrivons des simulations récentesde futurs débuts de période glaciaire réalisées avec l’EMIC de Potsdam, qui comporte une composante de cyclede carbone atmosphère-océan. L’une de ces simulations, dans laquelle 5 000 GtC sont libérées dansl’atmosphère par les activités humaines, mène à la conclusion que l’époque interglaciale actuelle durera aumoins un autre demi-million d’années, à cause de la capacité limitée des océans d’absorber une aussi importantelibération de carbone dans l’atmosphère.

that the method used by the EPICA Community Members toestimate the length of the interglacial MIS 11 is flawed, andthat a closer look at the δD data used by EPICA reveals thatthis interglacial probably lasted less than 10 kyr.) With regardto (2), several EMIC studies have shown that the long-termconcentration of atmospheric CO2 in the distant future is acrucial factor in determining when the next glacial inceptionmight occur (e.g., Berger and Loutre, 2002; Archer andGanopolski, 2005; Cochelin et al., 2006).

The remainder of this paper is structured as follows. A briefoverview of the Milankovitch theory of glacial inception isgiven in Section 2, and a review of some earlier work on sim-ulations of the LGI is presented in Section 3. In Section 4, theMcGill Paleoclimate Model (MPM) is described in the contextof other EMICs, and in Section 5 simulations of the LGI withboth the geophysical and ‘green’ MPM are given. Possible sce-narios for the next glacial inception are presented in Section 6,and some concluding remarks are given in Section 7.

2 On the theory of glacial inceptions The Milankovitch (1941) theory of long-term (multi-millen-nial) climate change states that a glacial inception occurswhen the summer insolation at high northern latitudesdecreases substantially and reaches a very low value.†

However, changes in the concentration of greenhouse gasesin the atmosphere amplify the glacial inception process,

through different feedbacks (e.g., Gallée et al., 1992;Ruddiman, 2003b).

The solar forcing at the top of the atmosphere slowlychanges due to variations of three orbital parameters of theEarth’s motion about the sun: (1) the eccentricity, with peri-odicities of around 100 and 400 kyr, (2) the obliquity, with aperiod of 41 kyr, and (3) the climatic precession (a measureof the Earth-sun distance during the summer solstice), with adominant period of 23 kyr (see Milankovitch (1941) and theexpanded explanation in Ruddiman (2001)). Milankovitchargued that the optimal conditions to enter a glaciation are ahigh climatic precession, a high eccentricity and a low obliq-uity (which collectively result in a low insolation in summerat high northern latitudes and a low seasonal contrast). Suchconditions occurred around 116 kyr BP (see point A inFig. 2), and as a consequence, the insolation in June at 62.5°Nreached the very low value of about 440 W m–2 (see the low-est arrow in the left-hand side of Fig. 3) and a glacial incep-tion occurred (Fig. 1).

The evolution of the orbital parameters over the next100 kyr does not replicate the above conditions for glacialinception because of the small eccentricity and weak varia-tions in the climatic precession (which is modulated by theeccentricity) over this period (see left-hand side of Fig. 2).Since the eccentricity is currently near the end of a 400-kyrcycle, its value will be small for the next 100 kyr, as will thedecreases in summer insolation at high northern latitudes (seearrows with question marks in Fig. 3). In fact, during the next100 kyr, Fig. 3 shows that the next weak insolation of about465 W m–2 will occur at 50 kyr AP (after present, defined here

Glacial Inceptions: Past and Future / 319


Age (kyr)

enth

icδ

O(‰

)1

8

Fig. 1 Time series of δ18 O measurements taken from benthic foraminifera in an ocean sediment core from Ocean Drilling Program (ODP) Site 980 (55º29’N,14º42’W, 2179 m) over the last 0.5 million years (adapted from Wang et al. (2002) by permission of the American Geophysical Union). Benthic datawere obtained from Cibicidoides wuellerstorfi. This record is a proxy for global ice volume (increasing downward). The last interglacial period extend-ed over the period 120–130 kyr BP (before present; here ‘present’ is defined as 1950), the Last Glacial Maximum (LGM) occurred at 21 kyr BP, andthe present interglacial period (the Holocene) started at around 10 kyr BP. Peak glacial periods also occurred at approximately 135 kyr BP, 250 kyr BP,340 kyr BP and 420 kyr BP, during MIS 6, 8, 10 and 12, respectively. MIS stands for the ‘marine isotope stage’.

†According to Martin Claussen (personal communication, 2006), this theoryis based on an earlier hypothesis published in Köppen and Wegener (1924).

MIS

as after 1950). This insolation value is significantly largerthan the low insolation values characteristic of the last fourglacial inceptions (see arrows on the left-hand side of Fig. 3).Since the future summer insolation variations at high northernlatitudes will have low amplitudes during the next 100 kyr, itis not obvious when the next glacial inception might occur —in 50 kyr, 100 kyr or later? Because of our knowledge of pastCO2 concentrations in the atmosphere and their relation to the100-kyr glacial-interglacial cycles (EPICA, 2004), it is high-ly likely that the future atmospheric CO2 level will play animportant role in determining the occurrence of the nextglacial.

Antarctic ice core records show that the atmospheric CO2level varied between approximately 180 and 280 ppm duringthe past half-million years (EPICA Community Members,2004), with low values of around 180 ppm occurring duringthe peak glacials, and high values of around 280 ppm occur-ring during the interglacials. Since the Industrial Revolution,which started in the mid-eighteenth century, the CO2 level hassteadily increased to its current level of more than 380 ppmtoday (2007), and it will likely continue to increase to two-to-three times this value in the next 100 yr (IPCC, 2007).

Recently, Archer and Ganopolski (2005) have argued thatsuch large CO2 concentrations will decrease very slowly overthe next 100 kyr (because of the slow ocean uptake of atmos-pheric CO2), and thus we shall experience an extremely longinterglacial. Later in this paper, some of the simulations ofArcher and Ganopolski (2005), along with those performed atMcGill by Cochelin et al. (2006), will be reviewed with theaim of estimating how long the present interglacial will last.

3 Review of past simulations of the last glacial inception

As mentioned previously, many of the earlier simulationattempts of the LGI (e.g., Royer et al., 1983; Rind et al., 1989;Oglesby, 1990) were made with seasonally varying atmos-pheric GCMs that were forced with insolation and radiative(atmospheric CO2) fields appropriate for the timeslice ataround 116 kyr BP. However, these simulations failed to pro-duce substantial snow buildup at high northern latitudes andhence the studies pointed towards the importance of includingamplifying feedbacks due to vegetation (e.g., Gallimore andKutzbach, 1996; de Noblet et al., 1996), the oceans (Dongand Valdes, 1995; Khodri et al., 2001), sea ice (Yoshimori et



Fig. 2 Past and future variations of the Earth’s orbital parameters (as calculated by Berger (1978)) that affect long-term climate changes: The climatic pre-cession (top), the obliquity (middle), and the eccentricity (bottom). Point A marks the beginning of the last glacial inception (LGI) during MIS 5d (seeFig. 1). Negative values on the time axis refer to times before 1950 and positive values refer to the future, after 1950.

al., 2002), and the polar surface energy balance (Vettorettiand Peltier, 2003). Further, the sensitivity of LGI simulationsto the initial size of the Greenland ice sheet has been investi-gated by Kubatzki et al. (2006).

Khodri et al. (2001), in a coupled atmosphere-ocean GCMtimeslice study of the LGI, showed that in their model oceanfeedbacks lead to a cooling of the high northern latitudes,along with an increase in the moisture transport to the polar



Fig. 3 Insolation at the top of the atmosphere (TOA) at 62.5°N in June, between 500 kyr BP (negative values) and 500 kyr AP (positive values), as calculat-ed by Berger (1978). (Note: In this paper AP means after 1950.) A ‘W’ marks a warm interglacial period, and an arrow (without a question mark) indi-cates the time of a past glacial inception. Starting from the left, the arrows correspond to MIS 11, 9, 7 and 5d shown in Fig. 1. The arrows with a questionmark indicate the possible time of a future glacial inception.

W m

–2

latitudes. The latter leads to an increased delivery of snow tonorthern latitudes and the steady buildup of snow at around70°N (Fig. 4). However, since the model of Khodri et al.(2001) is run for only 100 yr under the same seasonal cyclefor the insolation at 115 kyr BP, the actual buildup of icesheets over several millennia cannot be simulated. In reality,the LGI started at 119 kyr BP, and by 115 kyr BP substantialice was present. Therefore, ideally one should start an LGIrun at around 120 kyr BP and let it continue to 115 kyr BP andbeyond. To carry out such simulations, long-term transientruns of climate system models are necessary, and the appro-priate tool to carry out such runs is the EMIC, which general-ly includes most interactive components of the climatesystem.

Among the first EMICs which coupled most componentsof the climate system is that of Gallée et al. (1991, 1992). Inthe 1991 paper, a 2-D (latitude-height) model of the northernhemisphere atmosphere (zonally averaged) was developedand coupled to an ocean mixed-layer, sea ice and land surfacecomponents. In Gallée et al. (1992), this four-component cli-mate model was asynchronously coupled to a model of thethree main northern ice sheets and the accompanyingbedrock. This coupled climate-ice–sheet model is oftenreferred to as the 2-D Louvain-la-Neuve (LLN) EMIC.Starting at 120 kyr BP, with forcing consisting of the astro-nomically derived insolation (Berger, 1978) and atmosphericCO2 concentration obtained from the Vostok ice core(Barnola et al., 1987), the model was able to simulate therapid latitudinal growth of the North American and Eurasianice sheets from 120 to 110 kyr BP, as well as the last glacialmaximum ice-sheet volume at 19 kyr BP (see Fig. 7a inGallée et al., 1992). The last two glacial cycles were also sim-ulated with this model by Loutre and Berger (2000).

Wang and Mysak (2002; hereafter referred to as WM2002)were the first to carry out transient run simulations of the LGIusing a global five-component EMIC which included a zon-ally averaged latitude-depth model for the ocean’s thermoha-line circulation (THC) (Wright and Stocker, 1991). The otherfour components of this EMIC are the 2-D (latitude-longi-tude) dynamic ice-sheet model of Marshall and Clarke(1997), and the atmosphere, sea-ice, and land-surface modelsas described in Wang and Mysak (2000). Under Milankovitchforcing and Vostok-derived atmospheric CO2 for the period122 to 110 kyr BP, the model (called the geophysical MPM)was used to investigate the mechanisms involved in the LGI,including those associated with the THC in particular. Anovel result found by WM2002 is that during the buildup ofnorthern hemisphere ice sheets during the LGI, the THC alsointensified, which allowed for large moisture transports fromthe oceans to the continents at high northern latititudes. Thismodel will be described in more detail in Section 4, and theLGI simulations using the MPM will be presented inSection 5a.

The earlier GCM timeslice studies that illustrated theimportance of vegetation in ice-sheet growth during the LGI(e.g., de Noblet et al., 1996) motivated Crucifix and Loutre

(2002), Kageyama et al. (2004), Calov et al. (2005a, 2005b),Wang et al. (2005) and Kubatzki et al. (2006) to conduct var-ious EMIC transient LGI simulations that included a dynam-ic vegetation component. Crucifix and Loutre (2002) showedhow vegetation works in synergy with snow cover and sea iceto produce inception conditions. Through the vegetation-albe-do feedback mechanism (which is similar to the familiar ice-albedo feedback, except that boreal vegetation is a highlyenergy-absorbing surface, whereas ice is a highly reflectivesurface), Kageyama et al. (2004) showed that vegetationchanges over North America (the northern boreal forest grad-ually disappears before inception) amplify the insolation-induced cooling and initial ice-sheet buildup there. Thisprocess does not occur over Eurasia in their model becausethe initial climate there is warmer, and vegetation is fartheraway from the taiga-tundra threshold which must be crossedfor glacial inception to occur. Further, if vegetation is fixed atinterglacial conditions, inception does not occur anywhere(see Fig. 1b in Kageyama et al. (2004)). Interestingly enough,Kubatzki et al. (2006) found that the LGI does not occur forfixed Eemian vegetation, but does occur for fixed present-dayvegetation. Consistent with Kageyama et al. (2004), Wang etal. (2005) found that due to the vegetation-albedo feedback,large ice-sheet buildup during the LGI occurred over NorthAmerica. However, in contrast to Kageyama et al. (2004) butin agreement with Calov et al. (2005a), Wang et al. (2005)simulated small ice-sheet buildup over northwest Europe andeastern Siberia when vegetation is fully interactive in themodel.

Archer and Ganopolski (2005) have been able to success-fully simulate the past five 100-kyr ice age cycles using theconceptual model of Paillard (1998) in which the thresholdfor glacial inception was made to depend on the CO2 concen-tration as determined from a stability analysis (Calov andGanopolski, 2005) performed with the CLIMBER-2 EMICfrom Potsdam (Brovkin et al., 2002) coupled to the 3-D ther-momechanical ice-sheet model Simulation Code forPolythermal Ice Sheets (SICOPOLIS) (Greve, 1997) (see left-hand side of Fig. 5). These cycles were forced by the astro-nomically derived insolation and Vostok-derived CO2.However, to simulate possible future glaciations realistically,an atmosphere-ocean and seafloor carbon cycle model(Archer, 2005) was used to determine future atmosphericCO2 levels resulting from 300, 1000 and 5000 GtC anthro-pogenic releases to the atmosphere, neglecting natural carboncycle variability. The partitioning of a large release ofanthropogenic carbon between the atmosphere and theCaCO3-buffered oceans is such that, in the absence of naturalCO2 forcing, approximately 7% of the anthropogenic CO2remains in the atmosphere 100 kyr after the perturbation (seeupper right-hand panel of Fig. 5). Thus, it appears that thelong-term future levels of atmospheric CO2 will likely be wellabove the pre-industrial value of 280 ppm. The impact of thisresult on future glaciations will be discussed further inSection 6, where other scenarios for the next glacial, obtainedby the McGill Earth System Modelling Group, will bepresented.



4 The McGill Paleoclimate Model (MPM) The MPM is a 2.5-D EMIC that has interactive atmosphere-land-sea–ice-ocean-ice–sheet components; its origin can betraced back to the zonally averaged coupled atmosphere-ocean model of Stocker et al. (1992) that was developed atMcGill nearly 20 years ago. However, unlike the model ofStocker et al., the MPM includes a moisture balance modelfor the atmosphere, as well as land surface, ice-sheet and sea-ice components; further, all components are seasonally vary-ing and the atmosphere is forced by astronomically derivedinsolation (Berger, 1978) and Vostok-derived atmosphericCO2. In addition, the zonal wind stress is a specified forcingfield over the oceans. We call the MPM a 2.5-D modelbecause while it has a 2-D ocean component (for latitude anddepth), the other components are sectorially averaged acrosseach ocean and continent, the domains of which are shown inFig. 6. Note that the model domain only extends from 75°S to75°N; hence the Arctic Ocean, some northern parts of NorthAmerica and Eurasia, and the Antarctic continent have beenomitted. The consequences of these limitations on ice-sheetbuildup will be discussed in the next section.

The ice-sheet component in the MPM is the vertically inte-grated dynamic part of the 3-D model of Marshall and Clarke(1997). It has a latitude-longitude resolution of 0.5° by 0.5°.The atmospheric component is the energy-moisture balancemodel of Fanning and Weaver (1996), with an improvedwater vapour-temperature feedback (see WM2002). Further,as described in WM2002, the atmospheric variables (surfaceair temperature (SAT), surface specific humidity and precipi-tation) are downscaled to 5° by 5° in the region 30°N to 75°N.For details on the sea-ice and land-surface components, seeWang and Mysak (2000).

This ‘geophysical’ MPM was extended by Wang et al.(2005a) to include a new land surface scheme (LSS) with

vegetation dynamics; this new model has become known asthe ‘green’ MPM. The LSS is characterized by the followingimprovements over the original version in Wang and Mysak(2000): (1) parameterization of deciduous and evergreen treesby using the model’s climatology and the output of thedynamic global Vegetation COntinuous DEscription(VECODE) model (Brovkin et al., 2002) which determines ineach grid cell, the fractions covered by trees, grass and desert;(2) parameterization of tree leaf budburst and leaf drop; (3)parameterization of the seasonal cycle of the grass and treeleaf area indices; and (4) calculation of the land surface albe-do by using vegetation-related parameters, snow depth andthe model’s climatology. In addition, a systematic parameter-ization of the solar energy disposition (Wang et al., 2004) isnow included in the green MPM. The green MPM’s simula-tion of the present-day climate compared with that in the geo-physical MPM is much improved. In particular, the strongseasonality of terrestrial vegetation and the associated landsurface albedo variations are in good agreement with severalsatellite observations of these quantities. It is shown that withthe explicit representation of the vegetation-albedo feedbackin the model, slow millennial-scale climate changes duringthe Holocene are particularly well simulated (Wang et al.,2005b). We shall see in Section 5b that this feedback alsoplays an important role in the LGI.

A detailed discussion of EMICs is given in Claussen et al.(2002), where it is shown that this class of models lies in themiddle of the broad spectrum of climate models, whichextends from ‘conceptual’ (box) models (e.g., Paillard, 1998)to ‘comprehensive’ high-resolution coupled atmosphere-ocean GCMs (see Fig. 1 in Claussen et al.). Generally speak-ing, each EMIC falls into one of three categories: (1) a 2-D or2.5-D model like the MPM which is based on a zonally aver-aged ocean THC; (2) a hybrid model in which the ocean



Fig. 4 Monthly time series (40-yr record) of snow depth at a glaciation-sensitive region in northern Canada (70°N, 80°W), as calculated by Khodri et al. (2001)from two 100-yr runs in a coupled Atmosphere-Ocean General Circulation Model (A-OGCM). Time starts from January of year 50 for the control (pre-sent-day) experiment (solid line) and for the 115 kyr BP (glaciation) experiment (dot-dashed line), in which there is a reduced insolation at high north-ern latitudes in summer (see Fig. 3). The snow depth is very stable in the control run, with no snow in summer, whereas the 115 kyr BP run shows anincrease in snow depth associated with perennial snow cover beginning around year 66 (month 792).

component say, is a coarse-resolution 3-D GCM, that is cou-pled to a simpler (e.g., energy-moisture balance) atmosphericmodel; or (3) a coarse-resolution 3-D atmosphere-oceanmodel with sea-ice and land-surface representations in whichmany of the processes are simplified. In addition to theMPM, the Bern 2.5-D model, CLIMBER-2 (Potsdam) andMoBiDiC (Louvain-la-Neuve) fit into category (1), whereasthe University of Victoria model fits into category (2).Among the models in category (3) are those from theMassachusetts Institute of Technology and the Russian

Academy of Sciences, and also the EMIC EcBilt-CLIO (fromLouvain-la-Neuve). For more details on these models, seeTables 1 and 2 in Claussen et al. (2002). A perusal of thesetables indicates that these EMICs or their updates haveinteractive components for the complete Earth system,including the biosphere. Descriptions of the present-day versions of many EMICs, including some that were not included in Claussen et al. (2002), can be found on the following website: http://www.pik-potsdam.de/emics/toe_05-06-07.pdf



Fig. 5 Simulation by Archer and Ganopolski (2005), reproduced by permission of the American Geophysical Union, of the past five glacial-interglacial cycles(past 500 kyr) and of possible future interglacial-glacial cycles during the next 500 kyr. For the future runs, the CLIMBER-2/SICOPOLIS ice sheetmodel is coupled to an atmosphere-ocean and seafloor carbon cycle model (Archer, 2005). For the past, the model is driven by atmospheric CO2 derivedfrom the Vostok ice core (Petit et al., 1999) and insolation changes as calculated by Berger (1978) (see left-hand side of panels a and b). The greencurves represent the natural evolution of climate (see panels c and d), and the blue, orange and red curves (on the right-hand side) represent, respec-tively, the results for short-term anthropogenic releases into the atmosphere of 300, 1000 and 5000 GtC. (a) Past and future p CO2 of the atmosphere,according to, respectively, Petit et al. (1999) and the carbon cycle model of Archer (2005). (b) June insolation at 65°N normalized and expressed inunits of the standard deviation, σ. 1 σ equals about 20 W m–2 . The green, blue, orange and red lines are the values of the critical insolation, i0, thattriggers glacial inception. (c) The interglacial periods simulated by the model. (d) Global temperatures simulated by the model (green, blue, orange andred curves) and the past temperature as estimated for the Vostok ice core (black curve).

a

b

c

d

From these model descriptions, it is clear that EMICs dohave their limitations. For example, they cannot be used tostudy interannual climate variability associated with the ElNiño Southern Oscillation (ENSO); this requires a high-reso-lution atmosphere-ocean model, especially in the tropics.However, they are an excellent tool to carry out simulationsof many millennia of climate history and to investigate theinteractions of as many components of the Earth system aspossible in an efficient manner. Moreover, they can be usedto explore, quite thoroughly, the parameter space of a model.Thus, they are more suitable for assessing uncertainty, whichGCMs can do to a significantly lesser extent. Finally, fromlong transient runs of EMICs, we can identify interestingtimeslices in the evolution of climate that can later be thor-oughly investigated with GCMs.

5 Simulation of the LGI with the MPM a Simulations Without Vegetation During the initiation phase of the LGI, from 122 to 110 kyrBP, the northern North Atlantic, south of Iceland, was rela-tively warm at the surface (see Ruddiman and McIntyre(1979) and Fig. 7 (this paper), middle curve), likely becauseof an intensified THC (McManus et al., 2002). However, inthe Norwegian Sea farther north, the sea surface temperature

(SST) started to drop during MIS 5e at around 125 kyr BP andthe surface waters cooled by 3°–4ºC by 120 kyr (e.g., Cortijoet al., 1994). This cooling at around 120 kyr BP is consistentwith earlier findings of Kellogg (1980). As the far northernparts of the main continents started to cool during the firsthalf of this period due to reduced summer insolation (see Fig.3 and also Fig. 11a in Section 5b), any moisture transportedto these regions from the warm ocean around 50°–60ºNwould have fallen as snow and remained there year after year.Thus, as the climate continued to cool because of the orbitalforcing, the accumulated snow would have led to rapid icesheet growth through the ice-albedo feedback. Indeed, fromFig. 7 (lower curve) we can infer that by 110 kyr BP (duringMIS 5d), a substantial amount of land ice formed. Accordingto Lambeck and Chapell (2001), the global sea level haddropped about 50–70 m by this time, which is equivalent to anice volume in the range of 20–28 × 106 km3. Ice sheets with thisvolume would have become unstable near the margins of theNorth Atlantic, leading to large iceberg discharges and icerafted debris (IRD) deposited in the deep ocean at around107 kyr BP (see top curve in Fig. 7). Clearly, a measure of thesuccess of any simulation of the LGI is whether large ice sheetsin the volume range above can build up over the northern con-tinents during a 5–10 kyr period after 120 kyr BP.



longitude (degrees)

latit

ude

(deg

rees

)

Fig. 6 Land-sea configuration for the MPM (the yellow grids correspond to Greenland). The north-south resolution in the model is 5º latitude, except acrossthe equator, where it is 10º.

Table 1 shows the seven different runs that were made withthe MPM in WM2002 to investigate the relative roles of var-ious processes deemed to be important for early stages of theLGI and the subsequent ice sheet growth. In control run 1, thefully coupled MPM was integrated from 122 to 110 kyr BPwith the radiative forcing shown in Fig. 1 of WM2002. (Thisforcing is also shown as the first 12 kyr of the time series inFig. 11 in Section 5b.) Important model features included inthis run are the elevation cooling effect of orography and aparameterization for the freezing of rain at high latitudes andrefreezing of glacial meltwater (see WM2002 for details). Inthe other runs, 2 to 7, different constraints or effects weretaken out of the model, until finally only Milankovitch forc-ing was used to drive the atmosphere-ice–sheet model com-ponents.

Figures 8a, 8b and 8c show, respectively, the time series oftotal, North American and Eurasian ice volumes for all theruns listed in Table 1. Clearly, the total ice-sheet growth (in

red), starting around 119 kyr BP, is most rapid for run 1, withthe total volume reaching about 15 × 106 km3 (Fig. 8a), whichis about two-thirds of the observed value, as measured by sealevel drop (Lambeck and Chapell, 2001). While this underes-timation could be due to shortcomings of the model (e.g., theice-sheet freezing and refreezing parameterization, the coarseresolution), it is most likely due to the limited domain of themodel; north of 75°N there is no ice-sheet formation, in con-trast to what is likely to have happened in reality.

Fixing the freshwater flux into the ocean (run 2) or the SST(run 3) has a relatively small impact on the total ice volumegrowth, in agreement with Kageyama et al. (2004). However,neglecting the elevation effect of orography (run 4) or thefreezing/refreezing parameterization (run 5) has a majorimpact on the growth. This first effect was also investigatedby Kageyama et al. (2004), who found that it did not have acrucial effect on ice-sheet growth. Finally, in theMilankovitch run 7, the ice sheet growth is very small; this



NeogloboquNeogloboqupachydermayyyppachypachydeapachydermapachydermmpachydermaaaa

Cibicdoideswuellerstorfi

Fig. 7 Paleoceanographic data taken from ODP site 980 in the northeast North Atlantic (55° 29’N, 14° 42’W) (J.F. McManus, personal communication, 2002).Top curve: ice rafted debris (IRD); middle curve: proxy for SST as derived from δ18O measurements of planktonic foraminifera (Neogloboquadrinapachyderma s.); bottom curve: proxy for global ice volume (increasing downward) derived from δ18O measurements of benthic foraminifera(Cibicidoides wuellerstorfi). MIS stands for ‘marine isotope stage’.

run helps to explain why running atmosphere-only GCMs forthe LGI is generally not successful in simulating the LGI.

Comparison of Figs 8b and 8c reveals that, in most runs,the ice volume growth is similar for both continents, which isat variance with the general belief that the ice sheets werelarger over North America than over Eurasia during the lastglacial (e.g., Turon, 1984). In Section 5b, we shall show that

when an interactive vegetation component is included in theMPM, this discrepancy is removed.

Figure 9a shows that during early glaciation, the THC inthe control run is intensified until around 116 kyr BP; this isdue to high latitude ocean cooling and reduced freshwaterfluxes into the North Atlantic (Fig. 9b), which results in lessbuoyancy in the water column there and consequently



TABLE 1. Experimental design for simulations of the LGI with the MPM. For run 2, P – E + R into the ocean is prescribed. When the SST is prescribed (run3), the atmosphere-ocean heat and freshwater fluxes are fixed at their initial values. ‘Refreezing’ means freezing of rain and refreezing of meltwater(reproduced from WM2002 by permission of the American Geophysical Union).

Run Coupling CO2 Mountain Freezing/Refreezing

1. Control run Fully coupled Vostok yes yes2. Fixed freshwater flux P-E+R prescribed Vostok yes yes

(into the ocean)3. Fixed ocean SST prescribed Vostok yes yes4. No mountain Fully coupled Vostok no yes5. No freezing/refreezing Fully coupled Vostok yes no6. No mountain and no Fully coupled Vostok no no

freezing/refreezing7. Milankovitch only SST prescribed 280 ppm no no

Fig. 8 Simulated ice volume growths for the different runs described in Table 1: (a) total ice volume, (b) North American ice volume, and (c) Eurasian icevolume. (Reproduced from WM2002 by permission of the American Geophysical Union.)

enhanced North Atlantic Deepwater Formation (NADW).The strong THC (8 Sv above the interglacial maximum)increased the land-sea thermal contrast in the model andhence produced large moisture fluxes to the land, which isfavourable for rapid ice-sheet growth. For run 2, the surfacefreshwater flux into the ocean was fixed, and the THC inten-sity increase during the first 5.7 kyr is only 4 Sv. This leadsto a drop in SST there (compared to run 1) and extensive sea-ice formation in the NADW region just prior to 116 kyr BP.This results in a lower heat loss to the atmosphere and hencea drop in the THC intensity (Fig. 9a, green curve).

Figure 10 shows snapshots of ice sheet distributions overthe northern continents at three timeslices for the control run.At 120 kyr BP (Fig. 10a), ice first appears in the vicinity ofthe northern Laurentide, Scandinavian and Siberian ice-sheetregions. By 116 kyr BP (Fig. 10b), these ice sheets haveexpanded and new ice sheets have formed over Alaska andeastern Canada. By 110 kyr BP (Fig. 10c), thick ice sheets oforder 3 km have formed over Alaska, eastern Canada andnortheastern Europe. It is unlikely that such large sheetsformed over Alaska during the LGI (A. Dyke, personal com-munication, 2002; W.F. Ruddiman, presonal communication,2007), because of the large mountains in this region and theincreased transport of latent and sensible heat there under

116 kyr BP orbital forcing that produced summer snow melt(Vettoretti and Peltier, 2003). Moreover, in contrast to obser-vations (Turon, 1984), the model overestimates the volume ofice formed in Eurasia during the LGI. However, the simula-tion of ice in Siberia between 116 and 110 kyr BP is consis-tent with the oxygen isotope evidence of Siberian glaciationduring MIS 5d presented by Karabanov et al. (1998).

b Simulations With VegetationIn the green MPM, the dynamic vegetation model VECODE,developed for use in EMICs by Brovkin et al. (2002), hasbeen interactively coupled to the geophysical MPM in orderto incorporate the biogeophysical vegetation–albedo feed-back. In addition, as noted in Section 4, a new land surfacescheme has been introduced (Wang et al., 2005a).Furthermore, in contrast to WM2002, the Greenland ice sheetis now explicitly resolved, being located in the western half ofthe North Atlantic as part of the ice sheet model, but attachedto the North American continent (see yellow grid boxes inFig. 6) for the purpose of coupling it to the other componentsof the MPM.

The green MPM has been forced with variable insolation(Fig. 11a) and atmospheric CO2 concentration (Fig. 11b) forthe period from 122 to 80 kyr BP, which is 30 kyr longer than



Fig. 9 The maximum THC intensities (20-yr means) (a), and freshwater flux anomalies (20-yr means) integrated over 45°–75°N (b) in the North Atlantic forthe different runs described in Table 1. (Reproduced from WM2002 by permission of the American Geophysical Union.)

Flux

Ano

mal

ies

(Sv)

Max

imum

THC

Inte

nsity

(Sv)

the run in WM2002. In Wang et al. (2005) results are pre-sented for both a control run (in which the vegetation is inter-active globally) and a number of sensitivity runs withvegetation fixed in different regions (see Table 1 in Wang et

al. for details). Here, the focus will be mainly on describingthe control run. Figure 12 shows the time series for the totalice volume (heavy green line), and also the ice volumes overNorth America (NA) and Eurasia (EUR). At 110 kyr BP, the



longtitude (degrees)

Fig. 10 Ice sheet thickness distributions over North America and Eurasia for the control run 1 in Table 1 at 120 kyr BP (a), 116 kyr BP (b), and 110 kyr BP (c).(Reproduced from WM2002 by permission of the American Geophysical Union.)

Insolationin June at 62 5ºN Atmospheric C concentrationrationr derived from VostokVostokV dataOCOC 2O2O

(a) (b)

W m

–2

Fig. 11 (a) Insolation at a high northern latitude in summer, as calculated by Berger (1978), and (b) Vostok-derived atmospheric CO2 concentration taken fromBanola et al. (1999), between 122 and 80 kyr BP.

total ice volume is slightly smaller than that in the control runin the geophysical MPM (see Fig. 8a); however, now weobserve that with interactive vegetation in the model, the icesheet growth over North America is much larger than overEurasia, a feature which is believed to have happened in real-ity (Turon, 1984). Further, we note that there is a small ice-sheet buildup over Eurasia at 110 kyr BP, which was notobtained by Kageyama et al. (2004). At the end of the run, onthe other hand, the ice volumes over each continent are com-parable, with the North American ice volume being slightlylarger. Also, at 80 kyr BP, the total ice volume simulated isapproximately equal to the observed volume, as estimatedfrom sea level changes.

The time series for the sea-level equivalent ice volume overthe integration period (dashed red curve in Fig. 12) showslarge fluctuations which may be due to massive iceberg dis-charges, a feature not simulated in the green MPM because ofthe lack of ice-sheet thermodynamics. The first large inter-ruption in ice-sheet growth is presumably due to the substan-tial insolation increase after 115 kyr BP (see Fig. 11) andsubsequent basal ice melt and rapid ice sliding, which led tothe iceberg discharges.

Figure 13 shows the simulated ice thickness distributions atsix timeslices of the control run. By 116 kyr BP, permanent

ice has appeared in Alaska, and the Laurentide, Scandinavianand Siberian regions; however, in contrast to the case in thegeophysical MPM, there is much less ice in Eurasia (compareFigs 13b and 10b). Between 116 and 80 kyr BP, the ice sheetscontinue to expand longitudinally, towards the centre of thecontinents, as well as southward. At the end of the run(Fig. 13f), substantial ice sheets appear over Alaska, Canadaand northwestern Europe. According to evidence from glacialdeposits and striation patterns, the large simulated ice sheetover Alaska is unrealistic (W.F. Ruddiman, personal commu-nication, 2007)

Figure 14 illustrates the evolution of the northern tree anddesert fractions in the control run (red curves). The simulatedtree fraction in the control run follows closely the 23-kyr pre-cessional cycle for the insolation (Fig. 11 a); the desert frac-tion changes are opposite to these insolation variations.Figure 14 also shows the results for two sensitivity experi-ments: the evolution of the above fractions for either fixedSAT (green curves) or precipitation (blue curves). Upon not-ing the similarity between the blue curves (with active SATand fixed precipitation) and the red ones, we conclude thatnorth of 60°N latitude the tree and desert fraction changes aredriven predominantly by temperature changes. This isbecause the change in the number of growing degree days



Sea level-equivalent ice volume from Lambeck and Chappell (2001)

Total ice volume produced by the ‘green’ MPM

Ice

Vol

ume

(x 1

06km

3 )

Fig. 12 Ice volume growth simulated by the green MPM between 122 and 80 kyr BP. Heavy green curve: total ice volume over North America and Eurasia;light green curves: ice volume simulated over North America (NA) and Eurasia (EUR). Dotted red line: sea-level equivalent global ice volume esti-mated by Lambeck and Chappell (2001). (Reproduced from Cochelin (2004) by permission.)

drives the high latitude vegetation changes (Wang et al.,2005a). An increase in the insolation during the warm seasonin this region increases both the length and the temperature ofthis season, which favours tree growth.

Figure 15 shows, for the control run, the modelled treefraction distribution at three timeslices (122, 100 and 80 kyrBP). By comparing these plots with the simultaneous icesheet distributions (in Figs 13a, 13d and 13f), we note that the



longitude (degrees)longitude (degrees)

Fig. 13 Simulated ice-sheet thickness distributions in the control run of the green MPM at (a) 122 kyr BP, (b) 116 kyr BP, (c) 110 kyr PG, (d) 100 kyr BP, (e)90 kyr BP, and (f) 80 kyr BP. (Reproduced from Wang et al. (2005) by permission of the American Geophysical Union.)

Fig. 14 Tree fractions in the total land (including ice sheets) and (b) desert fractions in total land (including ice sheets), averaged between 60° and 75°N in thegreen MPM, for the control experiment (red), the experiment with fixed SAT (green) and the experiment with fixed precipitation (blue) in the vegeta-tion component. (Reproduced from Wang et al. (2005) by permission of the American Geophysical Union.)

(a) 122 kyr BP

(b) 116 kyr BP

(c) 110 kyr BP

(e) 90 kyr BP

(d) 100 kyr BP

(f) 80 kyr BP

trees progressively disappear and the ice sheets build up andgive way to desert in the high northern latitudes. The deserts(not shown) therefore expand to the places where the icesheets are located. At 80 kyr BP, there are almost no treesanywhere in North America between 60° and 75°N due to theexpansion of the ice sheets. The same is true in Eurasiabetween 65° and 75°N. In high northern latitudes, the treelinehas shifted southward by 5° to 10° over the course of the sim-ulation. In view of this, the tree fraction averaged between60° and 75°N (see red curve in Fig. 14a) has substantiallydecreased between 122 and 80 kyr BP, whereas the desertfraction has greatly increased in this region (see red curve inFig. 14b). These changes in vegetation and the associatedchanges in surface albedo contribute to the expansion of theice sheets, owing to the positive vegetation-albedo and ice-albedo feedbacks.

6 Simulation of the next (possible) glaciation

Loutre and Berger (2000) ran the 2-D LLN hemisphericmodel for the next 130 kyr under orbital forcing and variousconstant concentrations of CO2 ranging from 210 to 290 ppm.The red dotted ice volume curve in the bottom panel of Fig.16, taken from the summary paper by Berger and Loutre(2002), shows that a glacial inception would occur immedi-ately for the lowest value of CO2 in the above range.However, for a future ‘natural’ CO2 variation similar to thevariations of the past 130 kyr (as seen in the Vostok ice core),the present interglacial would last for at least another 50 kyr(see solid curve in bottom panel of Fig. 16). This is becauseof the small variations in the high northern latitude summerinsolation for the next 50 kyr (see middle panel of Fig. 16).Upon introducing a global warming scenario starting at time



longitude (degrees)

latit

ude

(deg

rees

)la

titud

e (d

egre

es)

latit

ude

(deg

rees

)

Fig. 15 Tree fractions for the entire model area in the control run of the green MPM at (a) 122 kyr BP, (b) 100 kyr BP, and (c) 80 kyr BP. Greenland is attachedto the North American continent for plotting purposes. (Reproduced from Wang et al. (2005) by permission of the American Geophysical Union.)

zero (1950) in Fig. 16, in which the CO2 rises linearly from280 to 750 ppm over the next 200 yr and then slowly dropsback down to the pre-industrial value of 280 ppm over thenext 800 yr, the Greenland ice sheet melts initially but thenreforms. The model in this scenario forecasts the end of thepresent interglacial also in about 50 kyr (see dashed red curvein the bottom panel of Fig. 16), which is at the same time asthe present interglacial would end in the ‘natural’ CO2 run(solid curve).

In Cochelin et al. (2006), experiments similar to those inLoutre and Berger (2000) were performed with the (global)green MPM, which was integrated forward in time (starting at1950) for the next 100 kyr under orbital forcing (Berger,1978) and a variety of CO2 scenarios. The first set of simula-tions was run under various constant atmospheric CO2 levels.In the second set of simulations, the CO2 level rapidlyincreases over the first 350 yr to 1200 ppm in year 2300 andthen slowly decreases over the next 850 yr until it stabilizesat various levels at 1.2 kyr AP (see inset in Fig. 20). For theremaining 98.8 kyr, the atmospheric CO2 level remains con-stant. This variation in CO2 represents the inclusion of a largeglobal warming episode superimposed on the constant CO2scenario. As will be discussed later in this section, theassumption that atmospheric CO2 will ultimately return (after

1 to 2 kyr) to around a pre-industrial level after a large inputof carbon into the atmosphere may not be realistic (Archer,2005). Accordingly, if larger levels of CO2 remain in theatmosphere for thousands of years, the current interglacialcould last a very long time (Archer and Ganopolski, 2005).

Figure 17 shows the time series of ice volume over NorthAmerica for the next 100 kyr in the first set of experiments.Figure 18 illustrates the evolution of the maximum intensityof the THC for these experiments, and Fig. 19 portrays thetree and desert fraction changes averaged over the high north-ern latitudes.

From Fig. 17, we observe that, depending on the CO2 level,there are three possible types of evolution for the ice volume:an imminent glacial inception, a glacial inception in 50 kyr,or no glacial inception during the next 100 kyr.Mathematically speaking, the climate system model passesthrough two thresholds for glaciation/no glacial inception asthe CO2 concentration is increased.

For CO2 levels less than or equal to 270 ppm, the climateenters into a glacial period quite quickly. This general resultis consistent with that of Berger and Loutre (2002; seeFig. 16). Ice starts to build up in the west of the northern lat-itude region of North America and then slowly expands east-ward and southward (figure not shown). We note that the



Fig. 16 Long-term variations of eccentricity (top), June insolation at 65°N (middle), and simulated northern hemisphere ice volume (increasing downward) (bot-tom) from 200 kyr BP to 130 kyr AP (the present is defined as 1950). For the future simulations, three CO2 scenarios were used: last glacial-interglacialvalues, a human-induced (global warming) concentration which peaks at 750 ppm at 200 yr AP, and a constant concentration of 210 ppm. These sce-narios produced the ice volumes indicated by the solid, red dashed and red dotted curves, respectively. Simulation results from Loutre and Berger(2000); eccentricity and insolation from Berger (1978).

Insolation (W m–2)

evolution of the ice volume increase is CO2 dependent, witha larger and more rapid buildup for lower CO2 levels.

It is of interest to note here that the results shown in Fig. 17are, in a certain sense, consistent with Ruddiman’s earlyanthropogenic hypothesis (summarized in Ruddiman (2007)).Were it not for the widespread agricultural activities in theearly Holocene (which resulted in the release of CO2 and CH4into the atmosphere), Ruddiman argues that due to the natur-al carbon cycle, the CO2 concentration in the atmospheretoday would be around 250 ppm and that a glacial shouldhave started, thus ending the current warm interglacial. Ourresults (Fig. 17) show that for CO2 levels in the range of240–270 ppm, a glacial is imminent. In other words, ourmodel predicts a glaciation now at the low CO2 levels pro-posed by Ruddiman.

For CO2 levels between 280 and 290 ppm, a glacial incep-tion is simulated in 50 kyr (see black and yellow lines in

Fig. 17); thus for CO2 somewhere between 270 and 280 ppm,the first threshold of a CO2 level for glaciation in 50 kyr iscrossed. We again note that the rate of ice sheet growth isCO2 dependent; there is a fairly linear increase in ice volumefor a concentration of 280 ppm. For the higher concentrationof 290 ppm, however, the green MPM first simulates a slowbuildup of ice for about 25 kyr, and then a more rapid buildup.For both levels of CO2, the ice sheet first builds up overnorthwestern Canada and then expands eastward and south-ward (figure not shown). The appearance of the LaurentideIce Sheet (LIS) is also CO2 dependent; the higher the CO2concentration, the later the LIS is formed.

For concentrations greater than or equal to 300 ppm, thereis no glacial inception for the next 100 kyr in the green MPM(see green line in Fig. 17). Thus, there is a second thresholdfor CO2 levels between 290 and 300 ppm for which no glacialinception occurs in the next 100 kyr.



Ice

Vol

ume

(×10

6km

3 )

Fig. 17 Ice volume growth simulated by the green MPM over North America for the next 100 kyr, with constant CO2 scenarios ranging from 240 ppm (bluecurve) to 300 ppm (green curve). Note: here and in Figs 18–22, AP in the time axis means after 1950. Reproduced with kind permission from SpringerScience and Business Media (Cochelin et al., 2006).

Since the current atmospheric CO2 concentration is alreadygreater than 380 ppm due to anthropogenic activities, theimminent glaciation simulated by the green MPM for con-centrations less than or equal to 270 ppm is not a likely sce-nario. However, the later glacial inceptions at 50 kyr AP,shown in Fig. 17 for the larger (constant) concentrations of280 and 290 ppm, might be possible if (1) a considerable por-tion of current anthropogenic carbon emissions could be arti-ficially sequestered, or (2) the oceans and terrestrial biospherecould absorb, over the next millennium, almost all the largepresent-day carbon emissions into the atmosphere. In bothscenarios, this would allow the CO2 level in the atmosphereto return to a pre-industrial value of around 280 ppm. As willbe discussed, this second scenario is unlikely to happen. As tothe likelihood of the first scenario, that is beyond the scope ofthis paper.

The evolution of the maximum THC strength for the sevenconstant CO2 runs (Fig. 18) shows an overall pattern similarto that for ice volume. For cases in which a glacial inception

occurs during the next 100 kyr, the long-term increase in thestrength of the THC is about 4 Sv, which is about half theincrease found in WM2002 for their glacial inception controlrun. It is also interesting to note, for the entire duration of the300 ppm CO2 run, the existence of a small-amplitude quasi-periodic oscillation with a period of about 20–25 kyr in theTHC strength, a signal which is presumably due to the pre-cessional component of the orbital forcing which affects theSST at high northern latitudes and hence the rate of NADWformation. Such oscillations are also seen in the low-CO2level runs, once the ice sheets have built up.

Since ice first appears at high northern latitudes during aglacial inception, it is expected that changes in the tree anddesert cover would be most noticeable at these high latitudes,where vegetation responds mainly to temperature (Fig. 14).Figures 19a and 19b illustrate the evolution of the tree anddesert fractions, averaged over 60° to 75°N. The tree and desertfractions averaged over 30° to 60°N (not shown) have similarevolutions to those seen in Fig. 19, but with smaller amplitudes.



Max

imum

TH

C I

nten

sity

(S

v)

Fig. 18 Maximum THC intensity simulated by the green MPM over North America for the next 100 kyr, with constant CO2 scenarios ranging from 240 ppm(blue curve) to 300 ppm (green curve). Reproduced with kind permission from Springer Science and Business Media (Cochelin et al., 2006).

At the start of the runs, we note that the higher the CO2concentration, the higher the tree fraction (Fig. 19a) and thelower the desert fraction (Fig. 19b). This is because the cli-mate warms as the CO2 level increases, and this induces anextension of the tree area and a compensating decrease in thedesert area. For the next 100 kyr, the tree fraction closely fol-lows the northern summer insolation variations (see Fig. 16,middle panel), while the desert fraction is opposite in phase tothese variations. After the buildup of ice sheets for each run(excluding the one for 300 ppm), we observe that there isalways a superimposed trend on the precessional oscillations:downward for the tree fractions and upward for the desertfractions. This occurs because as the ice sheets build up, theclimate cools down, which in turn induces a progressivedecrease in the tree area and an increase in the desert area.

In the second set of experiments with an initial globalwarming episode (see inset in Fig. 20), we again observe theexistence of a threshold for the suppression of glaciation. Forconcentrations of 290 ppm or less, there is a glacial inceptionat 50 kyr AP. For CO2 concentrations of 300 ppm or more, thegreen MPM does not simulate any glacial inception for thenext 100 kyr. Thus, the threshold value for the suppression ofglaciation is between 290 and 300 ppm and is presumablyvery close to the one obtained above in the runs without aglobal warming episode. Thus, the addition of an initial glob-al warming episode did not modify the CO2 level over whicha glacial inception will not occur.

These results are consistent with those of Archer andGanopolski (2005) who simulated a glacial inception at50 kyr AP (see the thick blue dashed line in the right panel ofFig. 5c) for a short-term (about 200 yr) anthropogenic releaseof 300 GtC into the atmosphere. In their atmosphere-oceanand sediment carbon cycle component coupled to theCLIMBER-2/SICOPOLIS ice sheet model, this input of car-bon results in a long-term atmospheric CO2 concentration ofjust under 300 ppm (see the thin blue line in the right panel ofFig 5a). On the other hand, for an anthropogenic input of1000 GtC, which results in a long-term atmospheric concen-tration of just over 300 ppm (see the orange curve in the rightpanel of Fig. 5a), there is no glacial inception until around130 kyr AP (see the thick orange dashed curve in Fig. 5c).Finally, we note from Fig. 5 that for a 5000 GtC release intothe atmosphere over the next few hundred years, an amountequivalent to all the known fossil fuels, there is no glacialinception for the next 500 kyr because the long-term CO2concentration in the atmosphere decreases very slowly overthis period, eventually asymptoting to about 400 ppm. Archer(2005) argues that the atmospheric concentration of CO2stays quite high for some time after a large carbon inputbecause of (1) the acidification of the ocean due to anthro-pogenic CO2 uptake, which increases the Revelle buffer fac-tor, and (2) the gradual warming of the ocean due to the CO2increase in the atmosphere, which reduces the solubility ofCO2 in the ocean. Both of these effects limit the amount of



Des

ert

frac

tion

Tre

e fr

actio

n

Fig. 19 Tree (a) and desert (b) fractions averaged between 60° and 75°N simulated by the green MPM for the next 100 kyr, with constant CO2 scenarios rang-ing from 240 ppm (blue curve) to 300 ppm (green curve). Reproduced with kind permission from Springer Science and Business Media (Cochelin etal., 2006).

CO2 that the oceans can absorb, and give rise to the long tailin the atmospheric CO2 decrease seen in the right panel ofFig. 5a. At 100 kyr, for example, 7% of the fossil fuel carbonwill still reside in the atmosphere. In view of these results,Archer and Ganopolski (2005) believe that the intensity andduration of the projected interglacial period will be longerthan those seen in the last 2.6 million years.

The Greenland ice volume in the green MPM varies onlyslightly over the next 100 kyr in the presence of a globalwarming episode. The green MPM does not simulate any sig-nificant melting of the Greenland ice sheet, which contrastswith the suggestion of Gregory et al. (2004) for a sustainedglobal warming of more than 2.7°C. Due to the increase inprecipitation at high latitudes, the MPM simulated a slightincrease in the Greenland ice volume during the first 1200years. It should be noted, however, that the MPM has toocoarse a resolution to resolve the ice sheet satisfactorily. Thismight explain why these results are quite different from those

of Loutre and Berger (2000), who obtained an almost com-plete melting of the Greenland ice sheet for their long-termsimulation which included a global warming episode.

The response of the THC with an initial global warmingepisode is shown in Fig. 21. The maximum strength of theTHC in each case at first rapidly decreases in response to therelatively large atmospheric CO2 increase and subsequentwarming of the climate (Fig. 21a); this response is also seenin many EMICs run under various CO2 increase scenarios(Petoukhov et al., 2005). The THC strength then rapidlyincreases until about 1200 yr AP, after which time it slowlydecreases. For a CO2 concentration of 280 ppm, the maxi-mum intensity finally returns close to its initial value afteraround 10 kyr AP. This is not the case for the other two CO2concentrations, where the strength at 10 kyr AP is about0.7 Sv (0.3 Sv) smaller than the initial value for the case of anequilibrium CO2 value of 300 ppm (290 ppm). The evolutionof the THC strength for the rest of the run (Fig. 21b) then



ntration for

Ice

Vol

ume

(×10

6km

3 )

Fig. 20 Ice volume growth simulated over North America by the green MPM for the next 100 kyr, with a global warming episode (shown in the inset) followedby a constant CO2 scenario of 280 ppm (blue curve), 290 ppm (red curve) and 300 ppm (green curve). Reproduced with kind permission from SpringerScience and Business Media (Cochelin et al., 2006).

follows the ice volume curve and increases after 50 kyr APfor CO2 levels of 280 and 290 ppm. This increase in the THCintensity is due to the decreased buoyancy in the northernNorth Atlantic, caused by the long-term cooling of the planetand the decreased runoff into the North Atlantic because ofthe buildup of ice sheets.

Figures 22a and 22b illustrate the evolution of the highnorthern latitude tree and desert fractions for the three CO2runs. We observe first a rapid increase (decrease) and then arapid decrease (increase) of the tree (desert) fractions over thefirst 1 kyr. Since tree growth varies with temperature, the treefraction time series follows the prescribed variations ofatmospheric CO2 and SAT (not shown), whereas the desertfraction time series varies inversely with these quantities. Therates of change for the tree and desert fractions are compara-ble to what have been simulated and observed for the pre-industrial Holocene (e.g., see Wang et al., 2005b), and thusthe rates of change shown here are deemed to be reasonable.

After these large-amplitude variations, the tree and desertfractions oscillate about a mean level with the precessionalperiod, similar to the patterns seen in Fig. 19. However, in thecases where glaciation occurs at 50 kyr AP, the tree fractionstend downward (Fig. 22a) and the desert fractions tendupward (Fig. 22b).

To summarize our results with the green MPM, we notethat the addition of the rapid initial change in CO2 forcing hastriggered modifications of the early state of the climate sys-tem (THC strength, high latitude SAT (see Cochelin et al.,2006), and vegetation cover). However, after a few millennia,the ice volume growth, maximum THC strength and vegeta-tion fraction simulated under the three CO2 levels with theinitial warming episode superimposed (Figs 20, 21 and 22) arequite similar to the ones simulated under constant CO2 levelswithout the initial warming episode (see Figs 17, 18 and 19 forthe CO2 cases of 280, 290 and 300 ppm). Thus, after a few mil-lennia, the climate system as modelled here has little memory



Maximum THC intensity for the next 10 kyr Maximum THC intensity for the next 100 kyr

Max

imum

TH

C in

tens

ity (

Sv)

Max

imum

TH

C in

tens

ity (

Sv)

Fig. 21 Maximum THC intensity simulated by the green MPM for the next 100 kyr with an initial global warming episode (see Fig. 20), followed by a con-stant CO2 scenario of 280 ppm (blue curve), 290 ppm (red curve) and 300 ppm (green curve). Reproduced with kind permission from Springer Scienceand Business Media (Cochelin et al., 2006).

of its initial or early conditions. As Archer (2005) has argued,this will not likely be the case when an interactive carbon cyclecomponent is included in the model. With ocean carbon chem-istry and biology included, there are important long-term bio-geochemical feedbacks to the atmosphere which will helpretain elevated levels of CO2 following a large carbon inputinto the atmosphere due to anthropogenic activity.

7 Concluding remarksIn this paper, we have shown that an appropriate and valuabletool for investigating the nature and causes of a glacialinception is the Earth system Model of IntermediateComplexity (EMIC). With such a model, it is possible tocarry out long transient runs in which all the components ofthe Earth system evolve with time and interact with eachother in response to temporally and spatially varyingMilankovitch forcing and prescribed radiative (atmosphericCO2) forcing. In particular, we have shown that during aglacial inception, several important feedbacks come into playin the buildup of large ice sheets, namely, the ice-albedo feed-back, the vegetation-albedo feedback and the orography-tem-perature feedback (the elevation cooling effect).

Based on the work of Berger and Loutre (from Louvain-la-

Neuve, Belgium), Archer and Ganopolski (from Chicago andPotsdam (Germany), respectively) and Cochelin et al. (2006),it appears that, in contrast to previous interglacials during thepast half-million years, the present interglacial may be ratherlong lasting (tens of thousands of years or more). This iscaused by (1) the special nature of the Milankovitch forcingpattern over the next 100 kyr, and (2) the relatively large con-centration of greenhouse gases in the atmosphere due tohuman activities. If one assumes that in the not-too-distantfuture, the concentration of CO2 in the atmosphere returns topre-industrial levels (in the range of 280 to 290 ppm) after aglobal warming episode, then the next glacial could start ataround 50 kyr AP (see Fig. 20). However, Archer andGanopolski (2005) suggest that it is more likely that for thenext 100 kyr or more, the atmospheric CO2 concentration willremain well above the pre-industrial level because of the lim-ited ability of the oceans to absorb a large short-term releaseof carbon into the atmosphere (due to the burning of allknown fossil fuels). Under this scenario, the present inter-glacial could last longer than 100 kyr — even as long as ahalf-million years (see right panel of Fig. 5c).

To give more credibility to the above results, it would be ofinterest to carry out glacial inception experiments with some



Tre

e F

ract

ion

Des

ert

Fra

ctio

n

Fig. 22 Tree (a) and desert (b) fractions averaged between 60° and 75°N simulated by the green MPM for the next 100 kyr with an initial global warmingepisode (see Fig. 20), followed by a constant CO2 scenario of 280 ppm (blue curve), 290 ppm (red curve) and 300 ppm (green curve). Reproduced withkind permission from Springer Science and Business Media (Cochelin et al., 2006).

of the newer EMICs (e.g., LOVECLIM-1.0, as described inDriesschaert et al. (2007)) under future greenhouse gas sce-narios which lie between those used in Archer andGanopolski (2005) and Cochelin et al. (2006). Such scenar-ios, for example, would consist of a doubling or tripling of thepresent CO2 concentration over the next 100 yr, and thenmaintaining an elevated concentration for one or two thou-sand years (see Fig. 1a in Driesschaert et al. (2007)).However, ideally, one would like to carry out long-termfuture climate simulations with EMICs in which there is afully interactive global carbon cycle which takes into accountthe carbon-climate feedback over the long term. In such sim-ulations, the land and ocean surface CO2 fluxes would beradiatively interactive with the atmosphere. A strategy forcarrying out such climate change experiments over the shortterm with coupled GCMs is given in Hibbard et al. (2007),and it is conceivable that this strategy could be suitably mod-ified for use in long-term simulations with EMICs.

AcknowledgementsThis paper is based on an updated version of the AlfredWegener Medal lecture given by the author at the European

Geosciences Union General Assembly in Vienna, in April2006 and also on a presentation at the EMIC Workshop, spon-sored by the Collège de France, in Paris, in May 2007. Theauthor wishes to thank Zhaomin Wang, Anne-SophieCochelin and Yi Wang for their fruitful collaborations in thepast. Also, the author is indebted to Katherine Knowland forher technical assistance in the preparation of this paper, andto Alexandra Jahn and Zav Kothavala for their helpful com-ments on an earlier version of this manuscript. The construc-tive reviews of Andrey Ganopolski and William Ruddimanwhich helped to improve this paper are much appreciated.

The support of a Natural Sciences and EngineeringResearch Council (NSERC) Discovery Grant and a CanadianFoundation for Climate and Atmospheric Sciences (CFCAS)Project Grant for this work is gratefully acknowledged.

Finally, on a personal note, I have been most fortunate tohave had, over the past four decades, the support and interestof my late father, Stephen Mysak, in the work of my researchgroups at the University of British Columbia (1967–86) andMcGill University (1986–present).



ReferencesARCHER, D. 2005. The fate of fossil fuel CO2 in geologic time. J. Geophys.

Res. 110: C09S05, doi:10.1029/2004JC002625.ARCHER, D. and A. GANOPOLSKI. 2005. A movable trigger: fossil fuel CO2

and the onset of the next glaciation. Geochem. Geophys. Geosys. 6:Q05003, doi:10.1029/2004GC000891.

BARNOLA, J.-M.; Y. RAYNAUD, Y.S. KOROTKEVICH and C. LORIUS. 1987.Vostok ice core provides 160,000-yr record of atmospheric CO2. Nature,329: 408–414.

BARNOLA, J.M.; D. RAYNAUD, C. LORIUS and N.I. BARKOV. 1999. HistoricalCO2 record from the Vostok ice core. In: Trends: A Compendium of Dataon Global Change. Carbon Dioxide Inf. Anal. Cent., Oak Ridge Natl.Lab., U.S. Dep. of Energy, Oak Ridge, Tenn.

BERGER, A. 1978. Long-term variations of daily insolation and quaternaryclimatic changes. J. Atmos. Sci. 35: 2362–2367.

BERGER, A. and M.F. LOUTRE. 2002. An exceptionally long interglacialahead? Science, 297: 1287–1288.

BROECKER, W.S. 1998. The end of the present interglacial: how and when?Quat. Sci. Rev. 17: 689–694.

BROVKIN, V.; J. BENDTSEN, M. CLAUSSEN, A. GANOPOLSKI, C. KUBATZKI, V.

PETOUKHOV and A. ANDREEV. 2002. Carbon cycle, vegetation, and cli-mate dynamics in the Holocene: Experiments with the CLIMBER-2model. Global Biogeochem. Cycles, 16(4): 1139, doi:10.1029/2001GB001662.

CALOV, R. and A. GANOPOLSKI. 2005. Multistability and hysteresis in the cli-mate-cryosphere system under orbital forcing. Geophy. Res. Lett. 32:L21717, doi:10.1029/2005GL024518.

CALOV, R.; A. GANOPOLSKI, V. PETOUKHOV, M. CLAUSSEN, V. BROVKIN and C.

KUBATZKI. 2005a. Transient simulation of the last glacial inception withan atmosphere-ocean-vegetation-ice sheet model. Part I: Glacial inceptionas a bifurcation in the climate system. Clim. Dyn. 24: 545–561,doi:10.007/s00382-005-0007-6.

CALOV, R.; A. GANOPOLSKI, V. PETOUKHOV, M. CLAUSSEN, V. BROVKIN and C.

KUBATZKI. 2005b. Transient simulation of the last glacial inception withan atmosphere-ocean-vegetation-ice sheet model. Part II: Sensitivity andfeedback analysis. Clim. Dyn. 24: 563–576, doi:10.007/s00382-005-0008.

CLAUSSEN, M.; L.A. MYSAK, A.J. WEAVER, M. CRUCIFIX, T. FICHEFET, M.-F.

LOUTRE, S.L. WEBER, J. ALCAMO, V.A. ALEXEEV, A. BERGER, R. CALOV, A.

GANOPOLSKI, H. GOOSSE, G. LOHMANN, F. LUNKEIT, I.I. MOKHOV, V.

PETOUKHOV, P. STONE and Z. WANG. 2002. Earth system models of inter-mediate complexity: closing the gap in the spectrum of climate systemmodels. Clim. Dyn. 18: 579–586.

COCHELIN, A.-S. 2004. Simulation of glacial inceptions with the “green”McGill Paleoclimate Model. MSc. Thesis, McGill University, Montreal.Available as C2GCR Report No. 2004-2.

COCHELIN, A.-S.; L.A. MYSAK and Z. WANG. 2006. Simulation of long-termfuture climate changes with the green McGill paleoclimate model: Thenext glacial inception. Climatic Change, 79: 381–401,doi:10.1007/s10584-006-9099-1.

CORTIJO, E.; J.C. DUPLESSY, L. LABEYRIE, H. LECLAIRE, J. DUPRAT and T.C.E.

VAN WEERING. 1994. Eemian cooling in the Norwegian Sea and NorthAtlantic Ocean preceding continental ice-sheet growth. Nature, 372:446–449.

CRUCIFIX, M. and M.F. LOUTRE. 2002. Transient simulations over the lastinterglacial period (126–115 kyr BP): Feedback and forcing analysis.Clim. Dyn. 19: 417–433.

de NOBLET, N.I.; I.C. PRENTICE, S. JOUSSAUME, D. TEXIER, A. BOTTA and A.

HAXELTINE. 1996. Possible role of atmosphere-biosphere interactions intriggering the last glaciation. Geophys. Res. Lett. 23: 3191–3194.

DONG, B. and P.J. VALDES. 1995. Sensitivity studies of northern hemisphereglaciation using an atmospheric general circulation model. J. Clim. 8:2471–2496.

DRIESSCHAERT, E.; T. FICHEFET, H. GOOSSE, P. HUYBRECHTS, I. JANSSENS, A.

MOUCHET, G. MUNHOVEN, V. BROVKIN and S.L. WEBER. 2007. Modelingthe influence of Greenland ice sheet melting on the Atlantic meridionaloverturning circulation during the next millennia. Geophys. Res. Lett. 34:L10707, doi:10.1029/2007GL029516.

EPICA COMMUNITY MEMBERS. 2004. Eight glacial cycles from an Antarcticice core. Nature, 429: 623–628.

FANNING, A.F. and A.J. WEAVER. 1996. An atmospheric energy-moisture bal-ance model: Climatology, interpentadal climate change, and coupling toan ocean general circulation model. J. Geophys. Res. 101: 15111–15128.

GALLÉE, H.; J.P. VAN YPERSELE, T. FICHEFET, C. TRICOT and A.L. BERGER.

1991. Simulation of the last glacial cycle by a coupled 2-D climate-icesheet model. Part 1: The climate model. J. Geophys. Res. 96:13139–13161.

GALLÉE, H.; J.P. VAN YPERSELE, T. FICHEFET, I. MARSIAT, C. TRICOT and A.L.

BERGER. 1992. Simulation of the last glacial cycle by a coupled 2-D cli-mate-ice sheet model. Part 2: Response to insolation and CO2. J.Geosphys. Res. 97: 15713–15740.

GALLIMORE, R.G. and J.E. KUTZBACH. 1996. Role of orbitally induced changesin tundra area in the onset of glaciation. Nature, 381: 503–505.

GREGORY, J.M.; P. HUYBRECHTS and S.C.B. RAPER. 2004. Threatened loss ofthe Greenland ice-sheet. Nature, 428: 616.

GREVE, R. 1997. Application of a polythermal three-dimensional ice sheetmodel to the Greenland ice sheet: Response to steady-state and transientscenarios. J. Clim. 10: 901–918.

HIBBARD, K.A.; G.A. MEEHL, P.M. COX and P. FRIEDLINGSTEIN. 2007. A strate-gy for climate change stabilization experiments. EOS, Trans. Am.Geophys. Union, 88(20): 217, 219, 221.

IPCC. 2007. Summary for Policymakers. In: Climate Change 2007: ThePhysical Science Basis. Contribution of Working Group I to the FourthAssessment Report of the Intergovernmental Panel on Climate Change. S.Solomon. D. Qin, M. Manning, Z. Chen, M. Marquis, K.B. Averyt, M.Tignor and H.L. Miller (Eds). Cambridge University Press, Cambridge,United Kingdom and New York, NY, USA, 18 pp.

KAGEYAMA, M.; S. CHARBIT, C. RITZ, M. KHODRI and G. RAMSTEIN. 2004.Quantifying ice-sheet feedbacks during the last glacial inception. Geophys.Res. Lett. 31: L24203, doi:10.1029/2004GL021339.

KARABANOV, E.B.; A.A. PROKOPENKO, D.F. WILLIAMS and S.M. COLMAN. 1998.Evidence from Lake Baikal for Siberian glaciation during oxygen-isotopesubstage 5d. Quat. Res. 50: 46–55.

KELLOGG, T.B. 1980. Paleoclimatology and paleo-oceanography of theNorwegian and Greenland seas: glacial-interglacial contrasts. Boreas, 9:115–137.

KHODRI, M.; Y. LECLAINCHE, G. RAMSTEIN, P. BRACONNOT, O. MARTI and E.

CORTIJO. 2001. Simulating the amplification of orbital forcing by oceanfeedbacks in the last glaciation. Nature, 410: 570–574.

KÖPPEN, W. and A. WEGENER. 1924. Die Klimate der geologischen Vorzeit.Borntraeger, Berlin, 255 pp.

KUBATZKI, C.; M. CLAUSSEN, R. CALOV and A. GANOPOLSKI. 2006. Sensitivityof the last glacial inception to initial and surface conditions. Clim. Dyn. 27:333–344, doi:10.1007/s00382-006-0136-6.

KUKLA, G.J.; R.K MATTHEWS and M.J. MITCHELL. 1972. Present interglacial:how and when will it end? Quat. Res. 2: 261–269.

LAMBECK, K. and J. CHAPPELL. 2001. Sea level change through the last glacialcycle. Science, 292: 679–686.

LOUTRE, M.F. and A. BERGER. 2000. Future climatic changes: are we enteringan exceptionally long interglacial? Climatic Change, 46: 61–90.

MARSHALL, S.J. and G.K.C. CLARKE. 1997. A continuum mixture model of icestream thermomechanics in the Laurentide Ice Sheet, 1. Theory. J.Geophys. Res. 102: 20599–20613.

MCMANUS, J.F.; D.W. OPPO, L.D. KEIGWIN, J.L. CULLEN and G.C. BOND. 2002.Thermohaline circulation and prolonged interglacial warmth in the NorthAtlantic. Quat. Res. 58: 17–21.

MILANKOVITCH, M. 1941. Kanon der Erdbestrahlung und seine Anwendungauf das Eiszeitproblem. Serbian Academy of Science, Belgrade. 633 pp.

OGLESBY, R.J. 1990. Sensitivity of glaciation to initial snow cover, CO2, snowalbedo, and oceanic roughness in the NCAR GCM. Clim. Dyn. 4: 219–235.

PAILLARD, D. 1998. The timing of Pleistocene glaciations from a simple mul-tiple-state climate model. Nature, 391: 378–381.

PETIT, J.R.; J. JOUZEL, D. RAYNAUD, N.I. BARKOV, J.-M. BARNOLA, I. BASILE, M.

BENDER, J. CHAPPELLAZ, M. DAVIS, G. DELAYGUE, M. DELMOTTE, V. M.

KOTLYAKOV, M. LEGRAND, V. Y. LIPENKOV, C. LORIUS, L. PÉPIN, C. RITZ, E.

SALTZMAN and M. STIEVENARD. 1999. Climate and atmospheric history ofthe past 420,000 years from the Vostok ice core, Antarctica. Nature, 399:429–436, doi:10.1038/20859.

PETOUKHOV, V.; M. CLAUSSEN, A. BERGER, M. CRUCIFIX, M. EBY, A.V. ELISEEV,

T. FICHEFET, A. GANOPOLSKI, H. GOOSSE, I. KAMENKOVICH, I. MOKHOV, M.

MONTOYA, L.A. MYSAK, A. SOKOLOV, P. STONE, Z. WANG and A.J. WEAVER.

2005. EMIC intercomparison project (EMIP - CO2): Comparative analysisof EMIC simulations of current climate and equilibrium and transientresponses to atmospheric CO2 doubling. Clim. Dyn. 25: 363–385, doi:10.1007/s00382-005-0042-3.

RIND, D.; D. PETEET and G. KUKLA. 1989. Can Milankovitch orbital variationsinitiate the growth of ice sheets in a general circulation model? J. Geophys.Res. 94: 12851–12871.

ROYER, J.F.; M. DEQUE and P. PESTIAUX. 1983. Orbital forcing of the inceptionof the Laurentide ice sheet? Nature, 304: 43–46.

RUDDIMAN, W.F. 2001. Earth’s climate: past and future. W.H. Feeman andCompany, New York, 465 pp.

RUDDIMAN, W.F. 2003a. The anthropogenic greenhouse era began thousandsof years ago. Climatic Change, 61: 261–293.

RUDDIMAN, W.F. 2003b. Orbital insolation, ice volume and greenhouse gases.Quat. Sci. Rev. 22: 1597–1629.

RUDDIMAN, W.F. 2007. The early anthropogenic hypothesis: challenges andresponses. Rev. Geophys. 45: RG4001, doi:10.1029/2006RG000207.

RUDDIMAN, W.F. and A. MCINTYRE. 1979. Warmth of the subpolar northAtlantic Ocean during northern hemisphere ice-sheet growth. Science,204: 173–175.

STOCKER, T.F.; D.G. WRIGHT and L.A. MYSAK. 1992. A zonally averaged, cou-pled ocean-atmosphere model for paleoclimate studies. J. Clim. 5:773–797.

TARASOV, L. and W.R. PELTIER. 1999. Impact of thermomechanical ice sheetcoupling on a model of the 100 kyr ice age cycle. J. Geophys. Res. 104:9517–9545.

TURON, J.-L. 1984. Direct land/sea correlations in the last interglacial com-plex. Nature, 309: 673–676.

VETORETTI, G. and W.R. PELTIER. 2003. Post-Eemian glacial inception. Part II:Elements of a cryospheric moisture pump. J. Clim. 16: 912–927.

WANG, Y.; L.A. MYSAK, Z. WANG and V. BROVKIN. 2005a. The greening of theMcGill Paleoclimate Model. Part I: Improved land surface scheme withvegetation dynamics. Clim. Dyn. 24: 469–480, doi:10.1007/s00382-004-0515-9.

WANG, Y.; L.A. MYSAK, Z. WANG and V. BROVKIN. 2005b. The greening of theMcGill Paleoclimate Model. Part II: Simulation of Holocene millennial-scale natural climate changes. Clim. Dyn. 24: 481–496, doi:10.1007/s00382-004-0516-8.

WANG, Z. and L.A. MYSAK. 2000. A simple coupled atmosphere-ocean-sea ice-land surface model for climate and paleoclimate studies. J. Clim. 13:1150–1172.

WANG, Z. and L.A. MYSAK. 2002. Simulation of the last glacial inception andrapid ice sheet growth in the McGill Paleoclimate Model. Geophys. Res.Lett. 29(23): 2102, doi:10.1029/2002GL015120.

WANG, Z.; L.A. MYSAK and J.F. MCMANUS. 2002. Response of the thermo-haline circulation to cold climates. Paleoceanogr. 17(1): 1006,doi:10.1029/2000PA000587.

WANG, Z.; R.M. HU, L.A. MYSAK, J.P. BLANCHET and J. FENG. 2004. A parame-terization of solar energy disposition in the climate system. ATMOSPHERE-

OCEAN, 42(2): 113–125.WANG, Z.; A.-S.B. COCHELIN, L.A. MYSAK and Y. WANG. 2005. Simulation of

the last glacial inception with the green McGill Paleoclimate Model.Geophys. Res. Lett. 32: L12705, doi:10.1029/2005GL023047.

WRIGHT, D.G. and T.F. STOCKER. 1991. A zonally averaged ocean model forthe thermohaline circulation. Part I: Model development and flow dynam-ics. J. Phys. Oceanogr. 21: 1713–1724.

YOSHIMORI, M.; M. C. READER, A. J. WEAVER and N. A. MCFARLANE. 2002. Onthe causes of glacial inception at 116 kaBP. Clim. Dyn. 18: 383–402.



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Glacial Inceptions: Past and Future