The 8200 yr BP cold event in stable isotope records from the North Atlantic region

Global and Planetary Change 79 (2011) 288–302

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Global and Planetary Change

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The 8200 yr BP cold event in stable isotope records from the North Atlantic region

Timothy J. Daley a,d,⁎, Elizabeth R. Thomas b, Jonathan A. Holmes c, F. Alayne Street-Perrott d,Mark R. Chapman e, Julia C. Tindall f, Paul J. Valdes f, Neil J. Loader d, James D. Marshall g, Eric W. Wolff b,Philip J. Hopley h, Tim Atkinson c, Keith E. Barber i, Elizabeth H. Fisher g, Iain Robertson d,Paul D.M. Hughes i, C. Neil Roberts a

a School of Geography, Earth and Environmental Sciences, University of Plymouth, Drake Circus, Plymouth, PL4 8AA, UKb British Antarctic Survey, Natural Environment Research Council, High Cross, Madingley Road, Cambridge, CB3 0ET, UKc Environmental Change Research Centre, Department of Geography, University College London, Pearson Building, Gower Street, London, WC1E 6BT, UKd Department of Geography, College of Science, Swansea University, Swansea, SA2 8PP, UKe School of Environmental Sciences, University of East Anglia, Norwich, NR4 7TJ, UKf School of Geographical Sciences, University of Bristol, University Road, Clifton, Bristol BS8 1SS, UKg Department of Earth and Ocean Sciences, University of Liverpool, Liverpool, L69 3GP, UKh Birkbeck College, University of London, Malet Street, London, WC1E 7HX, UKi School of Geography, University of Southampton, Southampton, SO17 1BJ, UK

⁎ Corresponding author at: School of Geography, EartUniversity of Plymouth, Drake Circus, Plymouth, PL4 8A

E-mail address: [email protected] (T.J. Dale

0921-8181/$ – see front matter © 2011 Elsevier B.V. Adoi:10.1016/j.gloplacha.2011.03.006

a b s t r a c t
a r t i c l e i n f o
Article history:Received 10 June 2010Received in revised form 21 March 2011Accepted 24 March 2011Available online 21 April 2011

Keywords:82008.2 ka BP eventpalaeoclimatestable water isotopesdata model comparisonNorth Atlantic

An abrupt cold event ca. 8200 cal.yr BP, is believed to have been caused by the catastrophic release of ice-dammed meltwater from Lake Agassiz and associated disruption of the Atlantic Meridional OverturningCirculation (AMOC). Previous reviews have highlighted both the “ideal” nature of the 8200 yr event as a targetfor numerical model validation and the likely geographical restriction of the ensuing cold event to the circum-North Atlantic region but have cited a lack of sufficiently resolved palaeoclimatic records to test thishypothesis. We review the current set of high-resolution stable isotope records from multiple archives (lake,bog, marine and ice cores) in the North Atlantic region for the period 9200–7400 yr BP (present=AD 1950).The isotopic values of terrestrial records are closely linked to isotopic values of palaeoprecipitation. All sitesprovided evidence for at least one centennial-scale anomaly (beginning ~8500–8250 yr BP) that exceededbackground variability. No evidence for spatial or temporal transgression of the isotope anomalies wasidentified, implying that a simultaneous climate signal was observed in the circum-North Atlantic region.Comparison with new simulations using the UK Hadley Centre model HadCM3, which was isotope-enabled tosimulate changes in the stable isotope composition of precipitation and forced by freshwater input (“hosing”)of 5 Sverdrups (Sv) (0.005 km3/s), for 1 yr, indicated agreement with the observed decrease in the amplitudeof the isotope anomaly with distance from the NWNorth Atlantic. The model-simulated duration of the event,however, was consistently shorter than that observed in palaeoclimatic records. A review of evidence forforcing additional to the catastrophic release of meltwater from Lake Agassiz (solar variability, sea-icefeedback and longer-term meltwater history) suggested that reduced solar output did not directly coincidewith the 8200 yr event, but that a more complex history of meltwater discharges and sea-ice feedback mayhave conditioned the AMOC for sustained climatic impact.

h and Environmental Sciences,A, UK. Tel.: +44 1752 584887.y).

ll rights reserved.

© 2011 Elsevier B.V. All rights reserved.

1. Introduction

1.1. Defining the target for data-model comparison

Understanding the ability of climate models to predict futurechanges in climate is central to reaching national and internationalagreement on climate policy. Oneway to evaluate model performanceis through “hind-casting”: the comparison of model simulations of

past climate with estimates of past variability in climatic andenvironmental variables. Evidence for past changes in these variablesis often and necessarily provided from proxy evidence, using dataderived from sedimentary archives and ice cores. Stable isotope dataserve as universal proxy indicators of past climatic variables that arerecorded in multiple archives and which can be directly compared,not only with each other, but with isotope-enabled climate modeloutput. The early Holocene cooling event, which occurred around8200 yr ago, provides a promising testing ground for data-modelcomparisons (Alley and Ágústsdóttir, 2005; Wiersma and Renssen,2006) for the following reasons. First, the cause of the event isbelieved to be known (Barber et al., 1999; Clarke et al., 2004), thereby

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289T.J. Daley et al. / Global and Planetary Change 79 (2011) 288–302

providing the evidence necessary to define the magnitude, durationand nature of the climate perturbation. Secondly, it is the highestmagnitude event of the Holocene in the northern mid- to highlatitudes and, therefore, of sufficient amplitude to make it a viabletarget for numerical modelling (LeGrande et al., 2006; LeGrande andSchmidt, 2008). Thirdly, it represents an abrupt cold oscillation duringthe Holocene, offering insight into the sensitivity of the atmosphere–ocean system that has relevance for understanding the potentialeffects of modern increases in glacial melting rates (Chen et al., 2006;Clark et al., 2008). Following recent reviews (Alley and Ágústsdóttir,2005; Rohling and Pälike, 2005), additional records with hightemporal resolution have recently been published from the NorthAtlantic region (Ellison et al., 2006; Marshall et al., 2007; Thomaset al., 2007; Kleiven et al., 2008; Daley et al., 2009; Domínguez-Villaret al., 2009; Thornalley et al., 2009) (Fig. 1), providing much neededclarification as to the timing, structure, duration, amplitude andgeographical variation of the event, as expressed in differentcomponents of the atmosphere–ocean system. Here we provide asynthesis of key stable isotope records for the period 9200–7400 yr BPfor the North Atlantic region (all dates hereafter are in cal. yr BP unlessotherwise stated), whichwe define as the ocean and its sediments andthe surrounding terrestrial seaboard extending to 2000 km inlandfrom the coastal boundary. We review terrestrial evidence forvariations in the isotopic composition of precipitation (δp, whichmay refer either to δ18O or to δD) during this period and itsimplications for changes in atmospheric circulation. These recordsare then compared with evidence for changes in surface and deepocean characteristics and records of variations in solar forcing aspotential drivers for the atmospheric anomaly. Finally, the palaeodataare compared with output from the isotope-enabled version of the UKHadley Centre General Circulation Model (GCM; HadCM3), yieldinginsight into the ability of HadCM3 to simulate the observed changes

Fig. 1. Locations of sites used in this synthesis; solid circles = lake sediment cores, solid squtriangle = peat core. Dark arrows indicate warm surface ocean currents: GS = Gulf Streamocean currents: EGC = East Greenland Current, LC = Labrador Current. Thick open arrow i

and therefore to increase understanding of the causes of the event.This synthesis presents the collated results of the UK NaturalEnvironmental Research Council (NERC)-funded Rapid ClimateChange Science Programme (RAPID) ISOMAP-UK project, whichfunded several of the new records and the code development forGCM simulations presented here.

1.2. The 8200 yr BP event

The 8200 yr BP event is believed to have been caused by theinjection of ~1014 m3 of impounded glacial meltwater into theLabrador Sea ~8470±300 yr BP by the breaching of ice-dammedproglacial Lake Agassiz–Ojibway (hereafter Lake Agassiz) (Barberet al., 1999; Clark et al., 2001; Renssen et al., 2001; Clarke et al., 2004;LeGrande et al., 2006), which formed at the southern margin of thewaning Laurentide Ice Sheet ~11.7 14C ka BP (Alley and Ágústsdóttir,2005). The volume and routing of cold, fresh meltwater is believed tohave been sufficient to disrupt the Atlantic Meridional OverturningCirculation (AMOC) and its associated heat transport (Renssen et al.,2001; Wiersma and Renssen, 2006), resulting in a widespreadregional or hemispheric centennial-scale cold event (Thomas et al.,2007). The discharge of this meltwater is believed to have beencatastrophic (~5 Sv, or 0.005 km3/s, for 0.5 yr) (Teller et al., 2002;Clarke et al., 2004) and sub-glacial (Lajeunesse and St-Onge, 2008),evidenced by amega-ripple sand bed and iceberg scour in Hudson Bay(Josenhans and Zevenhuizen, 1990; Lajeunesse and St-Onge, 2008),by a characteristic red bed in Hudson Strait (Kerwin, 1996), by thepresence of a double layer of turbidite-rich sediment dated ~8450 yrBP in the Labrador Sea (Hillaire-Marcel et al., 2007) and by thefreshening of oceanic waters at Cartwright Saddle on the LabradorShelf (Andrews et al., 1999; Keigwin et al., 2005). There is evidence tosuggest that the final drainage took place via a two stage draw-down

ares = ocean cores, solid triangles = ice cores, open squares = speleothem sites, open, NAD = North Atlantic Drift, IC = Irminger Current. Grey arrows indicate cool surfacendicates routing of final drainage of glacial Lake Agassiz.

290 T.J. Daley et al. / Global and Planetary Change 79 (2011) 288–302

(Clarke et al., 2004; Hillaire-Marcel et al., 2007), with an initialoutburst, completed within a year (Winsor et al., 2005), succeeded byprolonged routing of meltwater through the Hudson Strait (Clarkeet al., 2004).

The 8200 yr BP event was identified in the North Atlantic region inGreenland ice cores, observed in the oxygen isotope record of GISP2 as a~200 yr cold period during which temperatures dropped by anestimated 6±2 °C at Summit Station, Greenland (Johnsen et al., 1992;Dansgaard et al., 1993; Alley et al., 1997). Further examination of theδ15N values in N2 trapped in the ice suggested that themagnitude of thecooling was 5.4–11.7 °C, with a best estimate of 7.4 °C (Leuenbergeret al., 1999). Evidence for a contemporaneous isotopic event was laterdiscovered in the sediments of the Ammersee, Germany (vonGrafenstein et al., 1999), and Lake Rõuge, Estonia (Veski et al., 2004).A very high magnitude (8‰) δ18Ocalcite signal from a speleothem fromsouthwest Ireland (Baldini et al., 2002) was subsequently shown to bean analytical artefact (Baldini et al., 2007).

Stratigraphical changes in other palaeoclimate proxy data derivedfrom a range of archives have been associated with the broader period~8450–7800 yr BP, from sites distributed globally. Reviews by Alleyand Ágústsdóttir (2005) and Rohling and Pälike (2005) both pointedout that in the majority of locations beyond the North Atlantic regionand for the given temporal resolution of the data, the signals detectedwere of longer duration (multi-centennial) and relatively lowmagnitude, indicating a distinction between the regional responseto meltwater injection and global climatic anomalies. Delimitation ofthe event as a circum-North Atlantic phenomenon has, however, untilrecently been constrained by a lack of data with sufficient temporalresolution to permit synchronisation with the Greenland ice cores.

Whilst the outburst hypothesis is currently themost widely accepted,the meltwater event is set within the context of a contemporaneousreduction in the level of solar output, as reconstructed by variations in10Be and 14C flux (Bond et al., 2001; Vonmoos et al., 2006) and changes inthe extent of North Atlantic and Arctic sea ice (Bond et al., 2001; Rohlingand Pälike, 2005;Massé et al., 2008;Muller et al., 2009; Vare et al., 2009).Any atmospheric anomaly during the period 8450–8000 yr BP thereforewould have been influenced by a variety of driverswith differing controlsoperating at different spatial and temporal scales. It is also plausible thatvariations in atmospheric circulation had additional feedback effectsthrough redistribution of precipitation and consequent freshening ofoceanic moisture sources leading to expansion of sea ice (Duplessy et al.,1992; Clarke et al., 2009; Muller et al., 2009). Indeed, model simulationshave demonstrated that freshening of the North Atlantic via transport oflow-salinitynear-surfacewater fromHudsonStrait tohigher latitudes canbe sufficient to produce abrupt cooling associated with sea-ice feedbackon its own without a shut-down or substantial reduction of the AMOC(Clarke et al., 2009).

1.3. δp and its significance in palaeoclimatology

The isotopic composition of precipitation (δp) is a powerful tracerof atmospheric dynamics and the hydrological cycle (Cole et al., 1999;Araguás-Araguás et al., 2000). δp is controlled by isotopic fraction-ation processes during the movement of water molecules through thehydrological cycle (Dansgaard, 1964). Records of palaeo-δp, therefore,provide evidence for changes in the prevailing atmospheric circula-tion regime and the rain-out history of air masses. As such, however,they are rarely either simple palaeothermometers or palaeo-raingauges (Araguás-Araguás et al., 2000; Daley et al., 2010). Instead,these records reflect complex variability in the hydrological system,including in the stable isotope composition of the ocean surface, theseasonality of precipitation, the prevailing moisture source or mixtureof sources and potential evaporation effects as well as the conden-sation temperature of precipitation (Siegenthaler and Oeschger, 1980;Rozanski et al., 1992; Rozanski et al., 1993; Jouzel et al., 1997; Coleet al., 1999; Araguás-Araguás et al., 2000; Darling and Talbot, 2003).

Whilst there is good evidence that δp and temperature are closelyassociated at high latitudes (Jouzel et al., 1997; Leuenberger et al.,1999; Jouzel et al., 2000), the correlation becomes weaker at lowerlatitudes where other factors, especially rainfall amount, are moreimportant controls (Araguás-Araguás et al., 2000). It is nowrecognised that the application of simple linear relationships toconvert between palaeo-δp and single meteorological variablesoutside polar latitudes must be undertaken with caution (Schmidtet al., 2007) particularly in the mid-latitudes, where evidence for pastchanges in atmospheric circulation (Dean et al., 2002; Hammarlundet al., 2002; Magny et al., 2003; Mayewski et al., 2004; Magny et al.,2006) raises questions over the temporal constancy of any modernΔT-Δδp relationships (Daley et al., 2010).

Despite their complexity, records of palaeo-δp provide evidencefor the magnitude of isotopic change that is directly related tophysically-based fractionation mechanisms and that can be numer-ically modelled. Advances in numerical modelling have now made itpossible to “enable” GCMs to simulate changes in δp alongside moretraditional variables such as temperature and precipitation amount(Jouzel et al., 2000; LeGrande et al., 2006; Tindall et al., 2009),eliminating the need to convert palaeo-δp data to palaeo-temperaturewith the associated uncertainties for data-model comparison(Schmidt et al., 2007). With the advance in isotope modelling(Schmidt et al., 2007; LeGrande and Schmidt, 2008; Tindall et al.,2009), palaeo isotope data and GCM simulations are now directlycomparable. This approach provides greater confidence in data-modelintercomparisons.

1.4. Freshwater-hosing experiments with numerical models

Previous modelling studies of the 8200 yr BP event have beenperformed using models of various degrees of sophistication. One ofthe first modelling studies was performed by Renssen et al. (2001)using the ECBILT-CLIO model. The forcing for this study was4.67×1014 m3 (equivalent to 15 Sv over 1 yr) of freshwater addedto the Labrador Sea over a timespan ranging from 10 yr to 500 yr.Modelled results were most consistent with palaeo records when theforcing was added as 0.75 Sv over 20 yr. A subsequent experiment,however, with the same freshwater pulse but different model initialconditions led to a reduced AMOC that did not recover. Whilst theforcing used in the Renssen et al. (2001) study was reasonable givenunderstanding of the magnitude of the freshwater outburst at thetime, subsequent research (Clarke et al., 2004) suggests that thequantity of freshwater input involved in the final drainage of LakeAgassiz was much smaller (~one sixth of this value).

A more realistic forcing (2.6 Sv over 2 yr) was later applied to theCLIMBER model by Bauer et al. (2004). The experiment simulatedAMOC slow-down and recovery to its initial value within 20 yr, fallingshort of the centennial-scale event evidenced in palaeo records. Baueret al. (2004) were able to extend the duration of the event by addingadditional freshwater in the form of “white noise”. Meissner and Clark(2006) included this additional freshwater forcing in a moresystematic way by changing the routing of freshwater supplied bythe melting ice sheets in accordance with Licciardi et al. (1999).

Additional analysis of the Renssen et al. (2001) study usingECBILT-CLIO indicated close agreement with the distribution of proxydata evidence (Wiersma and Renssen, 2006). A progressively lowermagnitude for the atmospheric anomaly was observed with increas-ing distance from the NWNorth Atlantic. Increased atmospheric zonalflow was associated with an increased meridional temperaturegradient across the North Atlantic, due in part to the extension ofsea ice associated with reduced AMOC (Wiersma and Renssen, 2006).

The multi-millennial or multi-centennial simulations describedabove lack the complexity or resolution of the fully coupledatmosphere–ocean GCMs that are generally used for predictions offuture climate change. Increased complexity was more recently

Table 1Record metadata—core location, elevation, temporal resolution and proxy data type, with start and end dates for identified “primary” anomalies.

Site (inorder oflatitude)

Latitude Longitude Siteelevation(m.a.s.l.)

Archive type Event start date(to nearest 10 yr)

Event end date(to nearest 10 yr)

Duration Maximummagnitudeof oxygen isotopeanomaly(‰)

Temporalresolutionof data

Proxydata type

NGRIP (1) 75.10 −42.32 2917 Ice core 8250 8090 160 1.5 Annual/bi-annual δ18Op/δDpGRIP (2) 72.58 −37.64 3230 Ice core 8250 8090 160 2.5 Annual/bi-annual δ18Op/δDpDye 3 (3) 65.18 −43.83 2480 Ice core 8250 8090 160 2.0 Annual/bi-annual δ18Op/δDpLake Rouge(4)

57.73 26.90 114 Lake 8350 8150 200 0.4 Bi-decadal(12–25 yr)

δ18Ocalcite

MD99-2251(5)

57.46 −27.90 −2620 Ocean core 8410 8250 160 n/a (Sub) decadal(6–10 yr)

% N. pachyderma

MD03-2665(6)

57.45 −48.61 −3440 Ocean core 8400 8250 150 n/a Decadal/centennial

δ18Osea surfacewater/δ13C benthiccalcite

Pippikin Pot(7)

54.22 −2.51 326 Speleothem 8270 8110 160 1.5 Bi-annual δ18Op

Hawes Water(8)

54.18 −2.80 8 Lake 8450 8300 150 1.1 Sub-decadal(7.6 yr/sample)

δ18Ocalcite

White ScarCave (9)

54.17 −2.44 259 Speleothem 8230 8120 110 2.0 Bi-annual δ18Ocalcite

Lough Avolla(10)

53.01 −9.04 90 Lake 8500 8350 150 0.8 Bi-decadal(~20 yr)

δ18Ocalcite

Nordans PondBog (11)

49.15 −53.58 60 Bog 8400 8230 170 4.5 (Sub) decadal δ18Op

The Ammersee(12)

47.10 11.02 −80 Lake 8250 8090 160 1.0 Decadal (~11 yr) δ18Op

KaiteCave (13) 43.03 −3.65 860 Speleothem 8210 8140 70 1.5 Sub-decadal(~5 yr)

δ18Ocalcite


achieved through forced freshwater input or “hosing” of the fullycoupled and isotope-enabled atmosphere–ocean GISS GCM (LeGrandeet al., 2006; LeGrande and Schmidt, 2008). Results from thismodelweredirectly comparable with palaeo isotope data as this model has theability to simulate δ18Op and δ18Oseawater. Multi-millennial simulationswere not feasible given the complexity of the GISSmodel. LeGrande andSchmidt (2008) therefore focussed on the modelled response to arealistic flood from Lake Agassiz in preference to longer-term routingchanges or ocean preconditioning that are also believed to have beenimportant (Meissner and Clark, 2006). In a number of experimentswithrealistic freshwater forcing (corresponding to a meltwater pulse, MWP,input of between 0.394×1014 m3 and 3.154×1014 m3) they found amaximum reduction of between 25% and 60% in the AMOC, with theanomaly lasting ~180 yr. In most of their experiments, however, themaximumresponseoccurred twodecades after theMWPwith agradualreturn to the initial conditions.

2. High resolution records of centennial-scale change ~8200 BP

2.1. Inventory of terrestrial isotopic records from the North Atlantic region

Table 1 provides a summary of the metadata for the palaeo recordsfor which stable isotope data are available for the North Atlanticregion. These include data from ice, carbonate and cellulose from arange of archives. Whereas those from ice cores are direct measure-ments of the isotope composition of palaeo precipitation, those fromsediments provide only indirect estimates. In some cases, however, asdiscussed further below, the sedimentary isotope data can beconverted to quantitative estimates of δp.

Oxygen isotope data from GRIP (Thomas et al., 2007) and NGRIP(North Greenland Ice Core Project members, 2004) (Fig. 1) arepresented at 10 cm sampling resolution (Fig. 2). A composite record isalso shown. This is a stack of four cores (GRIP, GISP2, NGRIP andDYE3), all normalised to the GRIP mean and standard deviationplotted as averages over 27.5 cm depth increments (approximately2.5 yr resolution) (Thomas et al., 2007) (Fig. 2).

Only one oxygen isotope record is currently available from thepeat archive for the region (Daley et al., 2009). Core NDN02/1 wastaken in 2002 from Newfoundland (49.150°N, 53.583°W, 60 m a.s.l.;Fig. 1). Isotopic data are available at contiguous 1 cm depth resolutionfor the period 8450–8200 yr BP and 4 cm resolution for the widerperiod 8700–7800 yr BP. It has been demonstrated that the isotopiccomposition of preserved Sphagnum is directly related to the isotopiccomposition of precipitation via a constant cellulose-precipitationfractionation factor of 1.0274±0.001 (Daley et al., 2010).

Lake sediment isotopic data are available from several locations(Fig. 1). TheAmmersee inGermany (47.10°N, 11.02°E) (vonGrafensteinet al., 1999) provides a ~10,000 yr record of variability in δ18Ocalcite

(15,500–5000 yr BP) frombenthic ostracod-valve carbonate. Estimationof δ18Op from δ18Ocalcite values requires a correction to be made forchanges in the temperature atwhich the calcitewas precipitated and anestimation of the extent to which lake water δ18O values evolvedthrough the changing influence of evaporation over time. The authorsconverted the δ18Ocalcite data to estimates of δp by assuming a constantbottom-water temperature for the Holocene in this deep, pre-Alpinelake, and correcting for vital offsets in the ostracod carbonate.Evaporative enrichment of the lake was regarded as negligible.

LakeRõuge (local nameRõugeTõugjärv), located in southernEstonia(57°44′N, 26°54′E) (Veski et al., 2004), provides a 2300-yr (9500–7200 yr BP) record of oxygen isotope variations in bulk carbonate.Higher resolution (12–25 yr) intervals were measured between 8450and 8000 yr BP and 50 yr intervals in other parts of the core. δ18Ocalcite

values were not converted to estimates of δp, but the amplitude of thevariations is likely to be a reasonable estimate of the amplitude ofchange in δp during the 8200 yr BP event (Veski et al., 2004).

HawesWater, a hard-water lake (0.08 km2 in extent, 12 mdeep, 8 mabove sea level), in a Carboniferous Limestone catchment in northLancashire, NW England (54.18°N, 2.80°W), provides a sub-decadalearly tomid- Holocene record of δ18Ocalcite from fine-grained endogeniccarbonate. Contemporary calcite is shown to have precipitated, withinthe uncertainties in the data, in equilibrium with the epilimnion waterand summer temperatures of 17–19 °C (Marshall et al., 2007), implyingthat the palaeo δ18Ocalcite record reflects variations in the δ18O values of

B

D

F

H

J

L

N

Lo

ug

h A

volla

δδ18O

c (‰

VP

DB

)

-6.0-5.8-5.6-5.4-5.2-5.0

Lak

e R

ou

ge

δ18O

c (‰

VP

DB

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-11.8

-11.6

-11.4

-11.2

-11.0

ND

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Op

(‰ V

SM

OW

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-10-9-8-7-6-5

MD

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665

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-0.4-0.20.00.20.40.60.81.01.21.4

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99-2

251

% N

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hyd

erm

a (s

)

01234567

Haw

esw

ater

δ18O

c (‰

VP

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-6.6-6.4-6.2-6.0-5.8-5.6-5.4-5.2-5.0

MD

03-2

665

N. p

ach

yder

ma

(s)

δ18O

c (

‰ V

PD

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2.12.22.32.42.52.62.7

Am

mer

see

δ18O

p(‰

VS

MO

W)

-10.8-10.6-10.4-10.2-10.0-9.8-9.6-9.4

Calendar yrs before present (BP; present = AD1950)7400 7600 7800 8000 8200 8400 8600 8800 9000 9200

LV

5δ18

Oc

(‰ V

PD

B)

-8.5-8.0-7.5-7.0-6.5-6.0-5.5

GR

IP a

nd

NG

RIP

δ18O

ice

(‰)

-37.0-36.5-36.0-35.5-35.0-34.5-34.0-33.5-33.0

No

rmal

ised

Gre

enla

nd

Ice

core

sδ18

Oic

e (‰

)

-37.5-37.0-36.5-36.0-35.5-35.0-34.5-34.0-33.5

Pip

pik

in P

ot

Cav

eδ18

Oc

(‰ V

PD

B)

-5.8-5.6-5.4-5.2-5.0-4.8-4.6-4.4-4.2

Calendar yrs before present (BP; present = AD1950)

7400 7600 7800 8000 8200 8400 8600 8800 9000 9200

ΔTS

I (W

m-2

)

-0.8-0.6-0.4-0.20.00.20.40.6

Wh

ite

Sca

r C

ave

δ18O

c (‰

VP

DB

)

-7.0-6.5-6.0-5.5-5.0-4.5

A

C

E

I

G

K

M


image of Fig.�2


summer lake water and, by implication, precipitation. Chironomid-inferred temperature estimates provide tighter constraint on δp forHawes Water than for either Lake Rõuge or Lough Avolla (below).

A recent record from Lough Avolla, Ireland (53.01°N, 9.04°W;90 m a.s.l.; Holmes, Marshall, W., and Roberts, unpublished) coversthe period from about 14,800 to 1000 yr BP with a resolution of~160 yr for the interval from ~6700 to 1000 yr BP rising to ~20 yr forthe interval from 14,800 to 6700 yr BP. The δ18Ocalcite signal wasderived from fine-grained endogenic calcite, believed to haveprecipitated in isotopic equilibriumwith summer water temperaturesin the epilimnion of the lake. Hence, it was probably closely related toδp as in the case of Hawes Water (Marshall et al., 2007).

Three new speleothem records covering the timespan of interesthave recently become available. Two records from caves in theCarboniferous Limestone karst of the Yorkshire Dales, NW England,cover the period from 10,760 yr BP to 4325 yr BP at low resolution and8400 to 7950 yr BP at sub-decadal resolution (Hopley, P. J,, Richards,D. L., Marshall, J. D., Atkinson, T., Mattey, D., Vonhof, H. B., Smart, P. L.,unpublished). Pippikin Pot (52.22°N, 2.51°W; 326 m a.s.l.) is located6 km NW of White Scar Cave (54.174°N, 2.44°W; 259 m a.s.l.), and20 km to the east ofHawesWater (8 m a.s.l.; Fig. 1). Low resolution δ18Omeasurements on fluid inclusions from Pippikin Pot provide palaeo-temperature estimates and confirm that δ18Ocalcite is closely related toδp. Differences in the structure of the three δ18Ocalcite records from NWEngland (Pippikin Pot,White Scar Cave and the lake record fromHawesWater)may reflect subtle differences in the seasonal or temporalmixingof the δp signal at each site. The third speleothem record is from KaiteCave (43°2′N, 3°39′W, 860 m a.s.l.) in northern Spain (Domínguez-Villar et al., 2009) and provides a sub-decadal Holocene record, derivedfrom δ18Ocalcite analyses of a 1 m long stalagmite (LV5) (Domínguez-Villar et al., 2009). As with the records from NW England, variations inδ18Ocalcite are closely related to that in local δ18Op (Domínguez-Villaret al., 2008). Unlike the higher latitude sites, however, modernmeasurements of δ18Op indicate closer coupling with the amount ofprecipitation thanwith local temperatures,more similar to relationshipsobserved in the tropics (Araguás-Araguás et al., 2000; Jouzel et al., 2000).Currently, lower δ18Op values correspond to an increased amountof precipitation in northern Spain (Domínguez-Villar et al., 2008).

2.2. Records of change in the surface and deep ocean characteristics ofthe North Atlantic

The sedimentological record of surface and deep ocean changes iscomplex, largely on account of the differential capacity of cores to resolvedecadal to centennial scale events (de Vernal and Hillaire-Marcel, 2000;Keigwin et al., 2005). Our analysis primarily focuses on records withsufficiently high accumulation rates to detect the signal associated withthe drainage of glacial Lake Agassiz. Several cores from the NW NorthAtlantic and Labrador Sea confirm that the period 9500–8000 yr BP waspunctuated by a sequence ofmeltwater discharge events (Hillaire-Marcelet al., 2007), yet in most records evidence of concomitant changes insurface and deep ocean proxies remains elusive. Two recent studies havebeen published with decadal sampling resolution, providing proxyrecords that crucially link changes in surface and deep-ocean character-istics to the age of the final drainage of Lake Agassiz. These are derived

Fig. 2. Synthesis of decadal-scale terrestrial isotopic timeseries, associated oceanic records anregion arranged by latitude (most northerly records at the top). A: Variability in Total Solaminimum of the year 1986 (Wm−2) (Steinhilber et al., 2009). B: Greenland composite δ18Oi

cores GRIP (dashed line) and NGRIP1 (Rasmussen et al., 2006; Vinther et al., 2006). D: Variatiin percent Neogloboquadrina pachyderma (s) in ocean core MD99-2251, south of Iceland (EllisMD03-2665 from the Eirik Drift (Kleiven et al., 2008). G: Benthic Cibicidoides wuellerstorfi δannually-banded Speleothem WSC97-10-5, from White Scar Cave, (Hopley et al., unpublishEngland (Marshall et al., 2007). J: Sub-decadal variation in δ18Ocalcite from Speleothem YD0lake-sediment calcite from Lough Avolla, western Ireland (Holmes, Marshall, W. and RoberNewfoundland (Daley et al., 2009). M: δ18Oprecipitation reconstruction from Lake-sedimentδ18Ocalcite from speleothem LV5, northern Spain (Domínguez-Villar et al., 2009).

from core stations MD99-2251 on the Gardar Drift, southwest of Iceland(Ellison et al., 2006) and MD03-2665 on the Eirik Drift, in the southernLabrador Sea (Kleiven et al., 2008). The records from MD99-2251 coverthe period 9200–7400 yr BP with sub-decadal resolution (6–10 yrsampling interval). The records from MD03-2665 provide decadalresolution for much of their length, with the exception of a short phase~8350±50 yr BP for which centennial resolution data are available(Fig. 2).

3. Results and discussion

3.1. Defining the isotopic “event” in the North Atlantic region

The palaeodata series are presented in Figs. 2 and 3, with allrecords arranged by latitude and longitude, respectively, against theirassociated chronologies. Table 2 gives the reported start and end datesand associated uncertainties for the largest isotopic anomalies withinthe time period 8700–7800 yr BP in each of the records presentedhere.

All records provide evidence for at least one significant anomalycharacterised by a decrease in isotopic values in the period 8700–7800 yr BP that exceeds background variability and has characteristiccentennial-scale duration. With the exception of only one record(LV5), in which this anomaly lasted ~70 yr (Domínguez-Villar et al.,2009), all the others indicate an oscillation lasting ~110–200 yr with amedian duration of ~160 yr and a mean of 150±30(1σ)yr (Table 2).We define this oscillation as the “primary” anomaly. Previous studieshave used a statistical definition of the “event” (i.e. a N1σ change)(Marshall et al., 2007; Thomas et al., 2007). We recognise, however,that themagnitude of 1σwill be different for each palaeo record as thevalue depends on the temporal length and density of data and thedistribution of data points throughout the record. A 1σ change in onerecord, therefore, is not directly comparable with another if they arenot both equally data-rich (i.e. temporally-resolved). Consequently,we maintain the simpler definition used here, which also allows fordetection, where it exists, of spatial attenuation of the isotopic signal.

A clear geographical pattern is observed with regard to theamplitude of the primary anomalies relative to background variabilityin the records (Table 3; Fig. 2). The isotopic amplitude of theatmospheric anomaly (decrease in δp values) decreases with distancefrom the Labrador Sea. The largest amplitude atmospheric anomaly(4.5‰) is observed in the subfossil Sphagnum moss record fromNewfoundland (Daley et al., 2009). This core site is located closest tothe hypothesised melt water route from glacial Lake Agassiz. Themaximum amplitude of the atmospheric anomaly in Greenland isobserved in GRIP (~2.5‰), followed sequentially by Dye 3 (2‰) andNGRIP (1.5‰) (Table 3) (Thomas et al., 2007). Sites in NW Englandregister changes between 1.5 and 2.0‰. The smallest amplitudeoscillation (0.4‰) is reported from the site furthest from the NorthAtlantic, Lake Rõuge (Veski et al., 2004).

High resolution analysis of the δ18O records from the four Greenlandice cores (GRIP, GISP2, NGRIP, andDye 3) revealed variability ondecadaland shorter time scales with warm spikes interrupting the cooling andsubsequent warming at ~8220 and ~8160 yr BP (Thomas et al., 2007).The existence of these spikes defined a “central” event lasting 69 yr

d total solar irradiance (TSI) for 9200–7400 yr BP (BP=1950 AD) for the North Atlanticr Irradiance (TSI) relative to the value of the PMOD composite during the solar cyclece record (Thomas et al., 2007). C: Two synchronised δ18Oice records from Greenland iceon in lake-sediment δ18Ocalcite from Lake Rõuge, Estonia (Veski et al., 2004). E: Variationon et al., 2006). F: Variation in oxygen isotopic composition ofN. pachyderma (s) in core13C values from core MD03-2665 (Kleiven et al., 2008). H: Variation in δ18Ocalcite fromed). I: Oxygen isotopic variation in lake-sediment calcite from Hawes Water, northern1, from Pippikin Pot Cave (Hopley et al., unpublished). K: Oxygen isotopic variation ints, unpublished). L: δ18Oprecipitation reconstruction from Sphagnum peat core NDN02/1,calcite from the Ammersee, Germany (von Grafenstein et al., 1999). N: Variation in

Calendar yrs before present (BP; present = AD1950)

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Fig. 3. Synthesis of decadal-scale terrestrial isotopic timeseries and associated oceanic records for 9200–7400 yr BP (BP=1950 AD) for the North Atlantic region arranged bylongitude (westerly records at the top; for sources, see Fig. 2 caption).


Table 2Age uncertainties associated with the start and end envelopes of the primary anomalies in each record.

Site (inorder oflatitude)

Event start date(to nearest10 yr)

Ageuncertainty(± yr)

Start dateenvelope

Event end date(to nearest10 yr)

Ageuncertainty(± yr)

End dateenvelope

Duration Comments

NGRIP (1) 8250 47 8300–8200

8090 47 8140–8040

160 Dating errors based on assessment of annual layercounting error of 2% (Vinther et al., 2006)

GRIP (2) 8250 47 8300–8200

8090 47 8140–8040


Dye 3 (3) 8250 47 8300–8200

8090 47 8140–8040


Lake Rouge (4) 8350 80 8430–8220

8150 80 8230–8070

200 Dating errors based on palaeomagnetic correlationwith other dated core (Veski et al., 2004)

MD99-2251 (5) 8410 Asymmetrical 8450–8290

8250 Asymmetrical 8300–8140

160 Dating errors not including 400 yr ΔR (2σ range oncalibrated 14C date)

MD03-2665 (6) 8400 50 8450–8350

8250 50 8300–8200

150 dating errors not including 400 yr ΔR (1σ range oncalibrated 14C dates provided in Kleiven et al., 2008)

Pippikin Pot (7) 8270 30 8300–8240

8110 30 8140–8080

160 Dating errors based on analytical error in U/Thdeterminations

HawesWater (8)

8450 120 8570–8330

8300 120 8420–8180


White ScarCave (9)

8230 30 8260–8200

8120 30 8150–8090


Lough Avolla(10)



150 Dating errors based on calibrated range from 2σ rangeon 14 C determinations

NordansPond Bog (11)


8230 40 8270–8190

170 dating errors based on calibrated range from 2σ rangeon 14 C determinations

The Ammersee(12)

8250 n/a 8250–8250

8090 n/a 8090–8090


Kaite Cave (13) 8210 34 8244–8176

8140 34 8174–8106

70 Matches Greenland ~70-yr central event(Thomas et al., 2007)


within the 160 yr oscillation in Greenland (Thomas et al., 2007). It ispossible that the shorter ~70 yrdurationevent, observed in the annuallylaminated stalagmite (LV5) from northern Spain (Domínguez-Villaret al., 2009) and within the ~200 yr oscillation reported in the δ18Orecord from Lake Rõuge, Estonia (Veski et al., 2004) corresponds to this“central” isotopic event identified in the Greenland ice cores.

The primary event was recorded contemporaneously by surfaceocean proxies from marine cores MD99-2251 and MD03-2665(Fig. 1). In MD99-2251, there was a significant increase in theabundance of the cold water planktonic foraminifer Neogloboquadrinapachyderma (sinistral-coiling), consistent with a prominent south-eastward extension of polar surface waters ~8280 yr BP. In MD03-2665, this cooling is evident as an increase in the δ18Ocalcite ofN. pachyderma (sinistral-coiling) and is associated with 13C-depletion

Table 3Comparison of measured palaeo-isotope data for duration and magnitude of an identified penabled HadCM3 forced by freshwater hosing of 5 Sv (5×106 m3 s−1) over one year.

Site(in order of latitude)

Duration Modelled duration Maxof oxanom(‰)

NGRIP (1) 160 40 1.5GRIP (2) 160 40 2.5Dye 3 (3) 160 75 2.0Lake Rouge (4) 200 40 0.4Pippikin Pot (7) 160 30 1.5Hawes Water (8) 150 30 1.1White Scar Cave (9) 110 30 2.0Lough Avolla (10) 150 30 0.8Nordans Pond Bog (11) 170 40 4.5The Ammersee (12) 160 40 1.0Kaite Cave (13) 70 40 1.5Mean 150 40 1.7Stdev 34 13 1.1

in δ13Ccalcite of the benthic foraminifer Cibicidioides wuellerstorfi,which is indicative of a reduced flow of Lower North Atlantic DeepWater (LNADW).

Several of the records show an additional oscillation that isseparated from the primary anomaly by ~200 yr. The character of theadditional anomaly, however, is not consistent between these records.In core NDN02/1 fromNewfoundland, themagnitude of the additionalanomaly falls within the noise range of variability through the fullHolocene record. In speleothem LV5 from northern Spain, theadditional anomaly (observed as a ~10 yr, 2‰ excursion) precedesthe primary anomaly (observed as a maximum 1.5‰ negativeexcursion with duration ~70 yr) representing a highly anomalousvariation that is not detected elsewhere within the temporalresolution of the current set of stable isotope records. In ocean core

rimary atmospheric isotope anomaly with the ensemble of 10 simulations of isotope-

imum magnitudeygen isotopealy

Modelled magnitude ofisotope anomaly infirst decade(‰)

Modelled typicalmagnitude of isotopeanomaly(‰)

2.5 1.22.5 1.03.0 1.51.0 0.41.5 0.51.5 0.51.5 0.51.5 0.51.5 0.60.9 0.41.5 0.51.7 0.70.7 0.4

Age (cal. yr BP; BP= 1950)7900 8000 8100 8200 8300 8400 8500 8600 8700 8800

Sit

e re

fere

nce

nu

mb

er0

2

4

6

8

10

12

14

Fig. 4. Distribution of dates of initiation (closed circles) and completion (open circles)and their associated uncertainties (grey bars) for the primary anomaly identified instable isotope records from the North Atlantic region. Site reference numbers are givenin Table 1. Dates of initiation and completion are estimated from records as publishedwith the exception of NDN02/1 (re-analysed here—see Supplementary onlinematerial), Lough Avolla, Pippikin Pot and White Scar Cave (in preparation: as providedby the lead authors). Uncertainty ranges are specific to each record. For records 1, 2 and3, the grey bars represent the estimated uncertainty associated with annual layercounting error. For record 4, the grey bars represent uncertainty associated withpalaeomagnetic correlation with an additional core (Veski et al., 2004). For records 5and 11, the grey bars represent the calibrated range from the 2σ range on 14Cdeterminations. For record 6, the grey bars represent the calibrated range from the 1σrange on 14C determinations. Local reservoir corrections applied to both marine records(5 and 6) are consistent with those in the marine offset database (http://calib.qub.ac.uk/marine) but are not included in the grey bars. For records 7, 8, 9 and 13 the grey barsrepresent the analytical error on U/Th determinations. For record 10, grey barsrepresent the calibrated range on the 2σ range on 14C determinations on terrestrialmacrofossils. No information on uncertainty is available for record 12 because thepublished chronology was tied to GISP2 (von Grafenstein et al., 1999).


MD99-2251, a preceding additional anomaly was observed lasting~100 yr (8650–8550 yr BP) with a magnitude ~half that of theprimary anomaly. Given the inconsistent nature and detection ofthese additional anomalies, it is possible that they represent localphenomena, distinct from a clear circum-North Atlantic signaldetectable throughout the proxy records and characterised by anoceanic and atmospheric isotopic excursion lasting ~150±30(1σ)yr.

3.2. The timing of the event in different localities: time transgressivebehaviour or chronological uncertainty?

Despite similarities in the duration and relative amplitude of theprimary isotopic anomaly in each of the records, synchronisation ofthe event is restricted by differences in timing that are not currentlyresolvable given the dating uncertainties associated with the records(Fig. 4; Table 2).

Four records register the initiation of the primary anomaly ~8450–8400 yr BP. The published start date for the pronounced δ18O anomalyin NDN02/1 was 8450±90 yr BP, followed by an isotopic minimum atapproximately 8350 yr BP (Daley et al., 2009). Two minor agereversals in this core prompted re-analysis of the 23 published AMS14C dates using the iterative age-depth modelling program CLAM(Blaauw, 2010), yielding new start (~8400 yr BP) and end (~8230 yrBP) dates for this record, giving a duration of ~170 yr for the primaryevent and an isotopic minimum at ~8290 yr BP (Fig. 4; Table 2) (seeSupplementary online material for new age depth model andparameters used in CLAM). An age of 8380 yr BP was reported fromMD03-2665 on the Eirik Drift sediment (Kleiven et al., 2008) wheredeep water transport and sea surface temperatures declined almostcontemporaneously with the atmospheric signals recorded in NDN02/1and core HWLC1 from NW England (Fig. 4; Table 2). On Gardar Drift,

southwest of Iceland, a prolonged and non-linear reduction in deep flowspeeds occurred subsequent to considerable surface ocean fresheningcommensuratewith these other records. The remaining 7 sites have datesof initiation clustered ~8500–8250 yr BP (Table 2), although one of thesecan be excluded (vonGrafenstein et al., 1999) as its published chronologywas aligned with that of GISP2.

The similarity in the structure and duration of the primary event ineach of the records, and the apparent coincidence of a ~70 yr event inSpain with that reported from Greenland, imply that all the sites wereresponding to similar forcing. Therefore two options are possible toexplain the apparent age differences. Either, 1) the event wassynchronous across the North Atlantic region and the differencesbetween records result from dating inaccuracies; or 2) the ocean-atmospheric transmission of the signal operated via a mechanism thatled to a ~200 yr delay between two distinct groups of records. Wereoption 2) correct, it would be expected, given the relative amplitudeof the anomaly in each of the records, that the signal woulddemonstrate clear spatial/temporal transgression, i.e. a spatial patternto the two clusters of dated events, consistent with existingunderstanding of the mechanisms of atmosphere–ocean interaction.With the chronologies as provided, there is no evidence for a coherentspatial or temporal transgression of the primary isotopic signal (Figs. 2and 3), supporting the interpretation of the event as synchronousacross the North Atlantic region. Evidence for a reduction in theamplitude of the isotopic signal with distance from the hypothesisedsource of meltwater injection further supports the hypothesis that theevent was contemporaneous. For instance, a lagged response couldwell have operated through feedbacks (e.g. involving sea-ice extentand the distribution of atmospheric precipitation) that eitheramplified or damped the signal.

Option 1 would therefore seem more likely, but is difficult toreconcile within the constraints of the existing dating framework. Theseasonality of δ18O in precipitation is a valuable tool for dating icecores. At high accumulation sites, such as Dye 3 in southernGreenland, the annual layers can still be resolved back to the earlyHolocene. The Dye 3 oxygen isotope dated timescale (Hammer et al.,1986) is believed to only contain a 0.5% error back to 6900 b2k(6850 yr BP); however, layer thinning and diffusion increases thiserror rapidly before this time. A multi-parameter approach was usedto date ice cores beyond this point, based largely on the seasonaldeposition of chemical species found in the ice. This method has beenused to produce the Greenland Ice Core Chronology 2005 (GICC05)that dates the GRIP and NGRIP ice cores back to 60,000 yr BP(Andersen et al., 2006; Rasmussen et al., 2006; Vinther et al., 2006;Svensson et al., 2008), with an estimated error of ~830 yr for theoldest ice.

A prominent volcanic eruption that occurred during the 8200 yr BPcooling provides a convenient tephra marker with which to cross-correlate the Greenland ice cores. A distinctive double peak in theelectrical conductivity measurements (ECM) and high fluorideconcentrations occurred during the δ18O minimum and were visiblein all four cores (Thomas et al., 2007). This was used to synchroniserecords on the GICC05 time scale giving an age of ~8240 b2k(~8190 yr BP)±~50 yr according to the revised Dye 3 time scale(Vinther et al., 2006). The GICC05 timescale generally agreeswell withthe independently dated GISP2 age scale across the 8200 yr BP event,with GISP2 yielding dates that are ~40 yr older than the correspond-ing GICC05 ages.

A number of possible errors are associated with the GICC05 dating.These include: imperfect core stratigraphy, core loss during drilling/handling of the cores, data loss during sampling/measuring of thecores, insufficient measurement resolution and misinterpretation ofthe records (Alley et al., 1997). Synchronising multiple cores, usingvolcanic horizons, minimises these errors (Vinther et al., 2006). Themost significant error associated with the Holocene part of the GICC05age scale, however, is misinterpretation. Uncertainty in the dating,

http://calib.qub.ac.uk/marine

http://calib.qub.ac.uk/marine


such as years where a feature cannot be confirmed as an annual layerin all available records, has been given a maximum dating error of0.5 yr. That is to say half of the number of uncertain years issubsequently included in the final time scale, whilst the other half isdiscarded. Whilst the quoted age uncertainty for the GICC05 age scaleat 8200 yr might be questioned, it seems very unlikely that as many as200 yr have been missed in the layer counting procedure. Addition-ally, using a comparison of anomalies in 10Be in the GRIP ice core and14C in tree rings, Muscheler et al. (2004) recently estimated theisotopic minimum at GRIP to be at 8150 yr BP, with an uncertainty ofless than 20 yr: this lies within the uncertainty of GICC05.

Chronological frameworks for the terrestrial records were con-structed by two independent techniques: AMS 14C dating of organicremains and U/Th dating of calcite deposits. In core NDN02/1, theauthors reported two very small (~25 cal.yr) age reversals in thechronology at exactly the time frame of interest (Daley et al., 2009).Our re-analysis of the published 23 AMS 14C dates has yielded ayounger date (~60 yrs later; 8450–8350 yr BP) for the initiation of theevent with a revised duration of ~170 yr (see Supplementary onlinematerial).

The chronology of the Pippikin Pot,White Scar Cave and Kaite Cave(LV5) speleothems was based on U/Th age determinations. Thestandard error of the age models was ±30, ±30 and±34 yr,respectively (Table 2). Each of the three speleothem age modelsplaced the peak of the event at ~8150 yr BP, in agreement with theGreenland ice cores. The duration of the event in the White Scar Caveand Kaite Cave stalagmites was further constrained by annual laminaecounting to ~110 yr and ~70 yr respectively.

The source of chronological uncertainty in the U/Th-dated sequencefrom Hawes Water relates to the effect of detrital contamination(Marshall et al., 2007). Slight overestimation of the sedimentary age ismore plausible than underestimation. Consequently, the start and enddates of the event could reasonably be aligned by up to ~120 yr closer tothe present (Marshall et al., 2007), producing better alignment withGreenland and with the updated chronology of Newfoundland. In LakeRõuge, the 4020-yr long early Holocene part of the record is based on afloating, continuously counted, varve chronology with an estimated1–2% counting error, anchored by palaeomagnetic correlation and by 5AMS 14C dates from terrestrial macrofossils (Veski et al., 2004).Consequently, Lake Rõuge has a relatively wide age uncertainty(Table 2; Fig. 4). The widest age uncertainty, however, is associatedwith the chronology for LoughAvolla. This is based on six AMS 14C dates,five of which were determined from terrestrial macrofossils and onefrom bulk organic matter, yielding age uncertainties±~250 yr. In thebase of the sequence, the start of the Holocene is clearly recognisablefrom the isotope stratigraphy, lithostratigraphy and further constrainedthrough pollen correlation.

The development of chronological frameworks for marine corestratigraphy is universally affected by the requirement to estimate thelocal marine 14C reservoir age associated with the core site and timeframe of interest. For MD99-2251, a constant 400 yr marine correctionwas used (Ellison et al., 2006). The sample age for the largest excursionthat is equated to the 8200 yr BP event corresponds to a 2σ range of8190–8360 yr BP. If the surfacewaters were comparatively “old” duringthe 8200 yr BP event then a further reservoir correction would berequired resulting in a shift of the calendrical age estimates to youngerestimates (perhaps by ~100 yr). Whilst this scenario may be plausible,there is no evidence to justify an additional correction for the datingbeyond that which was applied by the authors for MD99-2251 (Ellisonet al., 2006). The earlier cooling in sea surface temperatures (SSTs),describedpreviously as the “secondary anomaly”,wasnot dateddirectlybutwasassumed tohavea similar ageuncertainty. For coreMD03-2665,the situation is more complicated. Kleiven et al. (2008) used a 400 yrcorrection throughout the core and an additional 128 yr correction toallow for an enhanced influence of the East Greenland Current on a fewsamples. Thepublished 14C agedeterminationsprovide a calendrical age

range of 8190–8380 yr for the entire interval of interest (Kleiven et al.,2008). These age estimates are reported with a 1σ error but are wellconstrained stratigraphically by dates before and after the 8200 yr BPevent.

Evidently, therefore, there is scope, within provided datinguncertainties, to align the characteristic primary anomalies observedin the multi-archive palaeo records more closely. Fig. 5 shows thisalignment graphically, with each of the terrestrial records synchro-nised by start date. A symmetrical U-shaped pattern of rapid decline(phase 1), new temporary equilibrium (phase 2) and rapid recovery(phase 3) is observed in the records from several sites (e.g. C, I, J, L, M,and N). In comparison, only one site, for which the data are multi-decadally resolved, shows evidence for a rapid decline followed by agradual recovery (D). In those sites in which a U-shaped pattern isexhibited, it is estimated that phase 2 involved a quasi-stable state ofatmospheric climate system lasting ~80–120 yr. Whilst we cannotexclude the possibility that there were real differences in timing,synchronicity of the records is currently the simplest and mostplausible interpretation and yields excellent agreement between therecords (Fig. 5).

3.3. Examination of internal and external forcing factors (feedbacks)

Understanding the forcing of the 8200 yr BP event is critical forevaluating numerical model performance. Several published recordshave indicated a pronounced offset between the timing of thecentennial-scale cold event centred on ~8200 yr BP and the proposedmeltwater trigger. The catastrophic outburst from the Laurentide icesheet has been dated at 8470±300 yr BP (Barber et al., 1999) based on14C-dated bivalves that were found above and below a preserved700 km long discharge deposit in Hudson Strait (Kerwin, 1996). Asystematic offset in 14C dates associated with local marine 14Creservoir ages allowed Barber et al. (1999) to establish a drainagechronology for Lake Agassiz that coincided (within dating uncertainty)with the primary isotope event in Greenland, thereby establishing apotential causal mechanism for the 8200 yr BP event. Whilst thecorrection brought the meltwater discharge and climate event moreplausibly together, there remained a ~200 yr delay between the bestestimate of meltwater release and the beginning of the centennial-scale cold event in Greenland. This also applies to the independently-dated records from northern Spain, SW Ireland, NW England andEstonia (Fig. 2). Release of an updated age-scale for the Greenland icecores (Rasmussen et al., 2006; Vinther et al., 2006) has not resolvedthis temporal mismatch between some records.

Some sedimentary and geomorphological evidence, however,suggests that the meltwater outburst did not occur as a single release(Teller et al., 2002; Clarke et al., 2004; Hillaire-Marcel et al., 2007) butrather as two discrete jökulhaups of ~1-yr duration, with dischargesof 3.6 Sv and 1.6 Sv, respectively (1 Sv=106 m3 s−1) (Teller et al.,2002; Clarke et al., 2004). These jökulhaups would have occurred viathe diversion of previous freshwater flow through the St. Lawrenceriver system to a sub-glacial passageway beneath the Laurentide IceSheet (LIS) and out through Hudson Bay (Lajeunesse and St-Onge,2008; Clarke et al., 2009). Originally it was postulated that these twoevents were separated by ~10–50 yr (Clarke et al., 2004) but theduration of the gap between jökulhaups is no longer completely clearfrom the literature. An initial first drainage ~8450±44 yr BP, followedby refilling of Lake Agassiz over a ~200 yr interval, has been suggestedon the basis of evidence from highly resolved records of relative sealevel change from peats in the Rhine–Meuse delta, near Rotterdam,The Netherlands (Hijma and Cohen, 2010). The relative sea level rise(2.11±0.89 m) recorded in these sediments and associated with thedrainage of Lake Agassiz was double the size of previous estimates(0.4–1.4 m) (Tornqvist et al., 2004; Kendall et al., 2008). A pulseddrainage and refilling, in line with the interval between two distinctepisodes of surface water freshening in the North Atlantic southwest

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of Iceland (Ellison et al., 2006) was considered the most likelyexplanation (Hijma and Cohen, 2010).

A ~200 yr delay between jökulhaups is certainly possible withrespect to the upper and lower end member dates associated with adouble-layer of carbonate-rich turbidites in the Labrador Sea(Hillaire-Marcel et al., 2007). It could also provide an explanationfor a delay between an initial pulse of Lake Agassiz meltwater and themaximum isotopic response, if a second pulse was required to induceAMOC slowdown, where the first was insufficient. This morecomplicated forcing scenario involving resealing of the ice dam andpartial refilling of Lake Agassiz prior to a second flood pulse has been

recognised as possible (Clarke et al., 2009) but has not yet beenexplored in GCM simulations.

The final drainage of Lake Agassiz occurred within the context ofmultiple releases of freshwater from the Laurentide Ice Sheet throughtheperiod 9500–7000 yr BP (Jansson andKleman, 2004;Hillaire-Marcelet al., 2007; Carlson et al., 2008). The recession of the LIS from 8400 to7000 yr BP led to the formation of a large number of glacial lakes (theLabrador–Ungava lakes) (Jansson and Kleman, 2004). These drained in~30 meltwater pulses, with fluxes individually exceeding 0.015 Sv intothe Labrador Sea, Ungava Bay and Hudson Bay (Jansson and Kleman,2004). Using landform-based reconstructions of the LIS decay pattern


and the associated glacial lake evolution, it was estimated that thesepulsed releases resulted in a total inflow of ~6000 km3 of freshwaterinto the North Atlantic, on top of that from Lake Agassiz (Clarke et al.,2004; Jansson and Kleman, 2004; Hillaire-Marcel et al., 2007; Carlsonet al., 2008). Additionally, the recession of the LIS opened a route forfreshwater runoff from the western Canadian Plains to the Atlantic(Carlson et al., 2009). It has been estimated that this routing change ledto a sustained increase in freshwater discharge throughHudson Strait of0.13±0.03 Sv (Carlsonet al., 2009). It has beensuggested that sustainedre-routing of this quantity of meltwater subsequent to catastrophicdischarge was both sufficient to cause a reduction in the AMOC and,indeed,was required, in order to produce the extended centennial-scaleisotopic and climatic anomalies observed in palaeoclimate records(Meissner and Clark, 2006; Carlson et al., 2009). This implies that it maynot be necessary for the timing of initial Lake Agassiz release to coincidedirectly with changes in oceanic and atmospheric proxies. Rather, itsuggests that the oceanic-atmospheric response to meltwater forcing iscontingent upon the preceding meltwater history and its cumulativeeffect on ocean salinity.

Solar variability is an additional potential candidate for forcingclimate change through variations in Earth's radiative balance.Variations in solar output, reconstructed from the 10Be content ofthe GRIP ice core, are available for the full Holocene (Muscheler et al.,2004; Vonmoos et al., 2006). 10Be production in the atmosphereoccurs as cosmic radiation interacts with nitrogen and oxygen. Theintensity of cosmic radiation is modulated by the strength of the solarmagnetic field. Variations in atmospheric 10Be therefore reflectvariations in the solar magnetic field, which are related to solaroutput (Steinhilber et al., 2009). 10Be deposition on Greenland is inpart influenced by atmospheric deposition processes, but anydistortion in reconstruction of the atmospheric values has beenshown to be much less than the variability due to changes in solaroutput (Vonmoos et al., 2006). Hence, the 10Be record can beconverted to Total Solar Irradiance (TSI; Wm−2; Fig. 2A) (Steinhilberet al., 2009). There is evidence for a reduction in solar output of~1 Wm−2 with a duration of ~150 yr, beginning ~8450 yr BP (Fig. 2A)(Steinhilber et al., 2009). The coincidence with similarly-timedvariations in some of the terrestrial and oceanic records presentedhere is striking. The TSI data were published on the GRIP ss09 ice coretimescale (Vonmoos et al., 2006; Steinhilber et al., 2009). Withcorrection of these data to the new GICC05 timescale, the isotopicevent in GRIP, recorded as a decrease in δp values, remainsasynchronous with the reduction in TSI, beginning after the recoveryin TSI. One cannot exclude completely the possibility that thereduction in solar output may have influenced climate in the NorthAtlantic region in advance of the ~150 yr primary isotopic eventsreported from Greenland and other records. We can safely assume,however, that reduced solar output did not coincide with the primaryisotopic anomaly in Greenland given that they are asynchronouswithin the same core (Thomas et al., 2007; Steinhilber et al., 2009).

Sea-ice feedback is a potential amplifier of any externally-drivenclimatic signal (Holland et al., 2006; Clarke et al., 2009), but itsreconstruction is notoriously difficult (Norgaard-Pedersen, 2009).Direct proxy records measuring changes in the export of materialentrained by pack ice and released onmelting are often discontinuous(Norgaard-Pedersen, 2009). The siliceous remains of diatoms specif-ically associated with activity in or near the ice are typicallysusceptible to dissolution and so are poorly preserved (Belt et al.,2007). A recent record from Fram Strait (Muller et al., 2009),however, has analysed variations in the biomarker IP25, which isbiosynthesised by Arctic sea-ice diatoms and serves as a proxy forpast seasonal spring sea ice (Belt et al., 2007; Massé et al., 2008;Muller et al., 2009; Vare et al., 2009). Preservation of this biomarkerhas been demonstrated for the last 30 kyr in Fram Strait (Muller et al.,2009). Both perennial sea ice and open waters equally result in nearabsence of the biomarker as neither set of conditions provides the

unique ecological niche required by the relevant species (Belt et al.,2007). Instead, diatom growth and therefore IP25 production is highwhen there is sufficient light penetration and where microscopiccapillaries and brine pockets exist, which typically occurs beneathfirst year ice (Belt et al., 2007). Combination with analysis of thephytoplankton-derived open-ocean proxy, brassicasterol, and organiccarbon accumulation, however, has yielded a 30 kyr-long record ofsea ice variability overlying the central conduit through which waterand sea ice exchanges between the Arctic and Atlantic Oceans(Aagaard and Carmack, 1989). The record from Fram Strait exhibitstwo spikes of increased IP25 flux at ~8200 yr BP and ~8000 yr BP,respectively. IP25 flux decreases for ~150–200 yr between thesespikes and is matched by a decrease in brassicasterol flux at a time ofrelatively high organic carbon accumulation (Muller et al., 2009). Theauthors suggest that the decrease in the fluxes of both IP25 andbrassicasterol probably reflects an abrupt increase in perennial sea iceof sufficient thickness and longevity to limit the growth of sea-icediatom assemblages. The spikes of increased IP25 flux both precedingand succeeding this trough represent transitional phases withseasonal sea ice in the surface waters overlying Fram Strait. It ispossible that the increase in carbon accumulation reflects enhancedsediment supply and faster burial resulting from increased glacialmeltwater activity through this period. Although similar biomarkerrecords for locations closer to our reviewed sites are not currentlyavailable, the records of N. pachyderma (s) abundance and isotopecomposition presented here (Fig. 2E–F) support a southwardextension of cold water masses and potentially support increasedsea ice cover given that the left-coiling (sinistral) variant has beendemonstrated to inhabit pack ice (Horner et al., 1992). Nonetheless,the record from Fram Strait provides a unique signal of increasedperennial sea ice in the region of East Greenland Current and WestSpitsbergen Current integration, a vital component of the AMOC(Aagaard and Carmack, 1989). In light of the simulated climaticinfluence of sea ice feedback from sea surface freshening alone(Clarke et al., 2009), this recorded increase in perennial sea-ice mayhave amplified considerably the climatic impact of the meltwaterforcing by Lake Agassiz drainage.

Evidence for climatic feedback resulting from a decrease in theAMOC was initially elusive (Keigwin et al., 2005) but is now providedby the high resolution oceanic records fromMD99-2251 (Ellison et al.,2006) and MD03-2665 (Kleiven et al., 2008) (Fig. 1). In MD99-2251, adecrease in the mean size of the sortable silt (SS) fraction indicated a~100–200 yr reduction in the flow speed of Iceland Scotland OverflowWater (ISOW), ~8260–8060 yr BP, before a rapid return to a fasterISOW current velocity, ~8000 yr BP (Ellison et al., 2006). ISOW is animportant component of the AMOC and combines with DenmarkStrait OverflowWater (DSOW) to form the returning deep water limbof the AMOC from the northern North Atlantic. In MD03-2665,decreased δ13C values of the benthic foraminifer Cibicidiodes wueller-storfi, reflected a ~150 yr period of reduced flow of Lower NorthAtlantic Deep Water (LNADW) that occurred simultaneously withsurface ocean cooling (Fig. 2), evident as an increase in the δ18Ocalcite

values of N. pachyderma (s) (Kleiven et al., 2008). The records fromboth of these cores provide evidence for coeval changes in surface anddeep ocean characteristics associated with the period of Lake Agassizfinal drainage.

The hypothesis that the primary isotopic anomaly was synchro-nous in the widely distributed, terrestrial and oceanic palaeoclimaterecords presented here is currently the most plausible, consistentwith the lack of clear evidence for time- or space- transgressivebehaviour andwith the existing dating control (Cheng et al., 2009). Bysynchronising the primary anomalies (Fig. 5), the current set ofisotopic records indicates that the event was ~150±30(1σ)yr. Theserecords still suggest that the event was related to the catastrophicdrainage of glacial Lake Agassiz, but that the pattern of meltwaterdischarge may have been more complicated than once believed,


possibly involving two major jökulhaups separated by up to 200 yr,superimposed on a background of: a) prolongedmeltwater re-routingthrough Hudson Bay; and b) many smaller jökulhaups from theLabrador–Ungava lakes.

3.4. Data–model comparisons

Simulations of the 8200 yr BP event have recently been performedwith a new version of the HadCM3 coupled atmosphere–ocean GCM(Gordon et al., 2000) which has been enabled with water isotopetracers as part of the ISOMAP-UK project (Tindall and Valdes, 2011-this issue). HadCM3 is a complex model and the main focus wason the response of the model to the final drainage of Lake Agassiz(no background meltwater flux, routing or re-routing was included;see Tindall and Valdes, 2011-this issue, for model parameters). Intotal, 10 different experiments were performed using HadCM3; ineach of these experiments, 5 Sv of freshwater were added to the NorthAtlantic for one year. The experiments varied in that, in each use, thefreshwater was added in an alternativemodel year corresponding to aslightly different initial state or over a different region of the NorthAtlantic. An ensemble of these experiments has been compiled toreduce the impact ofmodel “noise” or “weather” on the results, so thatthemodel's response to the freshwater forcing is more clearly defined.In reality some “weather” would have been associated with the8200 yr BP event, however, it is impossible to know the degree towhich this was simulated correctly (Tindall and Valdes, 2011-thisissue). Therefore, the ensemble mean is not necessarily the mostappropriate simulation to examine.

Table 3 shows the duration and amplitude of the isotope anomalyfor the ensemble of 10 experiments, for each of the terrestriallocations discussed in Section 3.1. In this case, the modelled durationof the event was defined as the length of time over which the isotopeanomaly remained more than 1σ below the control mean, analogousto the “central event” of Thomas et al. (2007). According to thisdefinition, the modelled event generally lasted ~40 yr, although atsome locations (Greenland and Newfoundland) δ18Op values fell togreater than 1σ below the mean in a secondary event. The modelledevent at Dye 3 appeared anomalous in that it lasted 75 yr, ~twice aslong as at other modelled locations. This apparent difference inbehaviour occurred because the event was stronger here than at otherGreenland locations and so the primary and secondary eventseffectively merged.

The modelled response of HadCM3 was strongest in the yearsimmediately following the MWP, consistent with other studies(LeGrande and Schmidt, 2008). Also shown in Table 3 is the typicalmodelled response for the first decade after the MWP and the typicalmodelled response for the event duration. There was good agreementbetween the magnitude of the modelled response for the first decadeand the maximum magnitude in palaeodata from nearly all sites,although the anomaly over Greenland was slightly larger in themodel. After the first decade, the modelled response weakenedconsiderably and its magnitude became typically one third to one halfthat suggested by the palaeodata.

The 8200 yr BP modelling results discussed here briefly and inmore detail by Tindall and Valdes (2011-this issue) support thefindings of previous modelling studies that sustained freshwaterforcing, in addition to that provided by a single flood from LakeAgassiz, is required to reproduce the observed structure and durationof the 8200 yr BP event (Bauer et al., 2004; Meissner and Clark, 2006).In each of the 10 model simulations reported here, the strongestresponse occurred in the decade immediately following the forcing,followed by a gradual return to the original climate. This contrasts,however, with the evidence for the duration and structure of theclimatic response from palaeo isotope data, where the event iscomparatively extended and, in most records, exhibits a U-shaped

structure with a rapid decrease, quasi-stability and rapid increase inδ18O values (Section 3.1; Fig. 5).

Comparison of measured isotope data from individual sites withthose simulated in grid-boxes representing hundreds of squarekilometres can be problematic given differences in the scale of spatialrepresentation. The model may simulate the correct general featuresbut with a small spatial error. Equally, palaeodata sites may be subjectto local forcing factors, irresolvable at the gridbox scale. Despite this,agreement is observed between HadCM3 output and the palaeorecords for the first decades of the simulation for most of the locationsdiscussed in this study. In Newfoundland, however, the model doesnot capture the full amplitude of the decrease in δp. Newfoundlandlies directly adjacent to the modern cold inner Labrador surfacecurrent, which makes this region particularly sensitive to upstreamchanges in this oceanic current and associated sea-ice variability. It ispossible that this site may have registered an amplified climaticanomaly through relatively localised sea-ice and atmospheric feed-backs, which cannot be accurately represented by the relatively coarseHadCM3 model grid.

Generally, the HadCM3modelling study (Tindall and Valdes, 2011-this issue) simulates the relative amplitude and spatial pattern of theprimary anomaly observed in each of the isotopic records throughoutthe North Atlantic region. The consistently short duration of themodelled event, however, implies that, in the real world, feedbackprocesses or a more complex pattern of meltwater release extendedthe length of the climatic impact of the freshwater perturbation andthat these processes were insufficiently well-represented in theHadCM3 experiments. It is unlikely that atmospheric processes alonewould have been sufficient to sustain the climate anomaly for thecentennial-scale duration observed from palaeo records, implying,therefore, that oceanic processes or prolonged meltwater influenceamplified the climatic anomaly.

4. Conclusions

We conclude that palaeo-isotope records frommultiple archives inthe North Atlantic region for the period 8700–7800 yr BP exhibit atleast one anomaly which exceeds background variability in eachrecord. Alignment of those “primary” anomalies is consistent with thelack of clear evidence for spatio-temporal transgressive behaviour andyields a regionally synchronous “event” that lasted 150±30(1σ)yr.The event probably occurred in response to a complex pattern ofmeltwater discharge associatedwith the final drainage of Lake Agassizand subsequent prolonged re-routing of meltwater that created asalinity anomaly that was sufficient for weakening AMOC. A “doublepulse” from Lake Agassiz may account for double spikes in several ofthe isotope records and also for the shorter duration of the isotopeminimum observed in the model experiments, which only involved asingle meltwater pulse as forcing. Despite a ~150 yr oscillation of~1 Wm−2 in total solar irradiance preceding the relative decrease inδp recorded in Greenland and elsewhere, we find no evidence fordirect causal linkage with variability in solar output.

New isotope-enabled GCM simulations broadly reproduce palaeo-isotope evidence for the spatial distribution of the climatic signal,showing that the amplitude of the maximum isotopic decrease inδ18Op declines with distance from the Labrador Sea. The duration ofthe modelled event, however, is consistently shorter than thatrecorded in the palaeodata from multiple archives. The isotopicanomaly in the real world may have been prolonged or amplified byadditional forcing or feedbacks, including sustained re-routing ofmeltwater and enhanced sea-ice feedback. Consistent with resultsfrom multi-centennial model simulations and with oceanic evidence,it implies that the oceanic-atmospheric response at 8200 yr BP to finaldrainage of Lake Agassiz was contingent upon the precedingmeltwater history, succeeding meltwater routing and its cumulative


effect on oceanic salinity.We recognise this proxy-model mismatch asa key challenge for future modelling experiments.

Acknowledgements

The authors acknowledge the support of NERC RAPID grant NER/T/S/2002/00460 to ISOMAP-UK and grant NE/C50907/1. We thankMilesIrving for drafting Fig. 1 and Dmitri Mauquoy for re-analysis of age-depthmodelling using radiocarbon age determinations. We thankWilMarshall for help with data from Lough Avolla. We also thank twoanonymous reviewers for their helpful comments in improving thismanuscript.

Appendix A. Supplementary data

Supplementary data to this article can be found online atdoi:10.1016/j.gloplacha.2011.03.006.

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The 8200 yr BP cold event in stable isotope records from the North Atlantic region

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