Battarbee R.W., Binney H.A. (Eds.) Natural Climate Variability and Global Warming: A Holocene Perspective

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The early to mid-Holocene

thermal optimum in the

North Atlantic

Eystein Jansen, Carin Andersson, Matthias Moros,

Kerim H. Nisancioglu, Birgitte F. Nyland

and Richard J. Telford

Keywords

Holocene, climate variability, ocean climate, North Atlantic, Nordic Seas

Introduction

The existence of an early to mid-Holocene thermal optimum is frequently referred

to in the paleoclimate science literature. (Early Holocene is here loosely deﬁned as

11–8 ka, and mid-Holocene is loosely deﬁned as 8–4 ka, both in terms of calendar

years BP.) The thermal optimum has been documented by a variety of paleocli-

matic evidence, spanning many Northern Hemisphere high-latitude paleoclimatic

archives, including ice cores, marine and limnic sediments, glacier records, and

speleothems. A comprehensive overview is given in the 4th Assessment Report of the

Intergovernmental Panel on Climate Change Working Group 1 (chapter 6, ﬁgure 6.9;

Jansen et al. 2007). Based on alkenone data, Kim et al. (2004) found that the

thermal optimum is most pronounced at high northern latitudes in the oceans,

and that in low-latitude areas it is less pronounced or nonexistent. In some low-

latitude areas, such as the Indian Ocean and parts of the tropical Atlantic, the

period of the thermal optimum (8–6 ka) as recorded in the high-latitude North

Atlantic is characterized by colder temperatures than in the late Holocene. Some

of the sites providing evidence for a colder early to mid-Holocene are from

upwelling regions; hence some of the cold anomalies may be related to enhanced

wind-driven upwelling. There appears to be a poleward ampliﬁcation of the ther-

mal anomaly, as well as a clear tendency that the warmest temperature interval

started earlier and lasted for a shorter period towards the Arctic than further south,

where it often appeared as a warm phase between about 8 to 6 ka (Koç et al. 1994;

Sarnthein et al. 2003; Duplessy et al. 2005; Hald et al. personal communication).

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Eystein Jansen et al.

Another feature is a delayed thermal maximum in the Labrador–Irminger Seas

south of Iceland, where peak Holocene temperatures occurred at about 6 ka. This

feature is probably related to the late demise of the Laurentide Ice Sheet (Andersen

et al. 2004; Hall et al. 2004; Berner 2006). Many authors have attributed the ther-

mal optimum to the combined effects of the demise of the continental ice sheets,

removing their inﬂuence on temperature distribution, and higher summer solar

irradiance than today due to orbital factors, especially the tilt of the Earth’s axis,

which provided a positive summertime radiative forcing anomaly in high-latitude

regions in both hemispheres.

Figure 5.1 shows selected localities where examples of the thermal optimum

have been demonstrated (Figure 5.2). Figure 5.3 summarizes the marine 6 ka

thermal anomaly in the Mediterranean and the north-east Atlantic–Nordic Seas

by using a data-set from alkenones only, thereby reducing errors when combining

different proxy methods. This ﬁgure also shows the northward ampliﬁcation of the

temperature anomaly.

Despite these ﬁndings, however, there are a number of North Atlantic and

Nordic Seas records where the thermal optimum is not apparent, and in some

there is instead a tendency for a warming throughout the Holocene

In this chapter we analyze these contrasting records in the high-latitude ocean

with the aim of attributing the contrasts to the forcing and/or to dynamics of the

ocean–atmosphere system, and we evaluate the interplay between the long-term

responses due to orbital forcing and shorter century to millennial scale variability.

–20

MD95-2011

Lake

Flarken

NEAP4K

Veste

Øykjamyrtjørn

OWS M

MD99-2269

Figure 5.1 Map showing

marine and terrestrial core

sites.

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The thermal optimum in the North Atlantic

125

Orbital forcing and climate model response

The effect of orbital forcing on sea-surface temperatures (SSTs) at 6 ka compared

with pre-industrial conditions has been studied in the Paleoclimate Modeling

Intercomparison Project PMIP2 with coupled ocean–atmosphere climate models

(AOGCMs) (see Cruciﬁx this volume, Figure 4.3). Similar simulations have also

been performed with other models, including models of intermediate complexity

(e.g. Liu et al. 2003; Rimbu et al. 2004; Renssen et al. 2005). A consistent feature

of these simulations is the positive summertime SST anomaly in the North

Atlantic–Nordic Seas area. The SST response of the MIT-EMIC (Michigan

Institute of Technology – Earth System Model of Intermediate Complexity) to the

6 ka orbital forcing compared with the response of the model to pre-industrial

forcing, i.e. the response solely due to changing orbital parameters, is shown in

Figure 5.4. The response is strongest in high latitudes, with very little response in

the tropics, and the most pronounced warming occurs near the Polar Front in the

North Atlantic–Nordic Seas. This poleward ampliﬁcation in the Atlantic sector is

consistent with observational evidence (Figure 5.3; data from Kim et al. 2004; Kim

and Schneider 2004a,b). The degree to which the models are able to separate

surface and thermocline temperatures and the seasonal response is variable. The

12 000 10 000 8000 6000 4000 2000 0

(a)

(b)

Lake Flarken

Vestre Øykjamyrtjørn

)C°( .pmet naem launnA

)C°( .pmet yluJ naeM

e years BP

(c)

)C°( TSS

MD95-2011

Figure 5.2 Examples of

records displaying a mid-

Holocene thermal maximum.

(a) UK37 data from MD95-

2011 (Calvo et al. 2002). (b)

Pollen-based annual mean

temperature reconstruction

from Lake Flarken (Seppä

et al. 2005). (c) Pollen-based

reconstruction of mean July

temperature from Vestre

Øykjamyrtjørn (Folgefonna

peninsula; Bjune et al. 2005).

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126

Eystein Jansen et al.

imprint of the seasonal forcing on the ocean sub-surface is clearly different

between models and observations, as noted in the following section. This implies

that for climate models to represent the true ocean sub-surface and thermocline

variability in response to the forcing, a better representation of high-latitude

surface ocean stratiﬁcation is required.

Ocean records and responses to orbital forcing

Sea-surface-temperature reconstructions

Reconstructions of surface layer temperatures of the high-latitude North Atlantic

and Nordic Seas through the Holocene have been obtained from various proxy

methods. Transfer functions have been produced based on data from alkenones,

–40

–20

20 20

–3

–2

–1

Temp. (°C)

Figure 5.3 6–0 ka alkenone

sea-surface temperature

differences for selected North

Atlantic and Mediterranean

cores. (Data from Kim and

Schneider 2004a,b.)

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The thermal optimum in the North Atlantic

127

coccolith species, diatom, and forminifers, as well as planktonic foraminiferal

oxygen isotope records and some Mg/Ca records, for parts of the Holocene.

Multiple proxies have been analyzed in some cores, enabling a direct comparison

of the output from the different methods without the errors of temporal correla-

tion. Core MD95-2011, obtained from beneath the main ﬂow of warm Atlantic

water towards the Arctic, is a good example (Figure 5.5). The Holocene tempera-

ture development reconstructed by the different proxy methods is very different,

both in terms of the amplitude of millennial and shorter time-scale variability, and

for overall Holocene trends. The proxies cluster into two distinct categories: those

–2

Insolation (W m

–2

)

–80 –60 –40 –20 0 20 40 60 80

–0.5

0.5

1.5

Latitude

–80 –60 –40 –20 0 20 40 60 80

Latitude

Temperature (°C)

Longitude

Latitude

–250 –200 –150 –100 –50 0 50 100

–50

–5

(b)

(a)

Figure 5.4 Annual mean

response in a simulation with

the Massachusetts Institute of

Technology Earth System

model of intermediate

complexity (Dutkiewicz et al.

2005) using 6 ka boundary

conditions. (a) 6 ka sea-

surface temperature (SST)

increase compared with pre-

industrial (ad 1850) SSTs.

(b) Annual mean insolation

change relative to pre-

industrial (upper panel), and

annual mean SST change

relative to pre-industrial.

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128

Eystein Jansen et al.

with a distinct Holocene thermal maximum (diatom-based SSTs, alkenones);

and foraminiferal-based proxies (transfer functions, species distribution, oxygen

isotope data corrected for ice-volume effects), which do not show an early to

mid-Holocene maximum. Instead, these proxies show the opposite, a warming

trend towards the late Holocene and more high-amplitude changes at century–

millennial time-scales after about 4 ka. The close correspondence between the

(a)

(b)

TSS

suguAt

C°( )

100

3.5

3.0

2.5

2.0

1.5

1.0

12 000 10 000 8000 6000 4000 2000 0

BDPV‰( O)

amredyhcap .N)nis(

.Namredyhcap)%( )nis(

e years BP

C°( TSS )sedoillub .G δ

‰)( O

2.0

1.8

1.6

1.4

1.2(c)

MD95-2011

NEAP4K

(d)

(e)

MD95-2011

Figure 5.5 Comparison

between sea-surface

temperature reconstructions

from: (a) alkenones (Calvo

et al. 2002); (b) diatoms in

MD95-2011 (Andersen et al.

2004; Koc personal

communication); (c)

planktonic stable oxygen

isotope data from NEAP4K

(Hall et al. 2004); (d) stable

oxygen isotope data and (e)

planktonic census data from

MD95-2011 (Andersson et al.

2003; Risebrobakken et al.

2003).

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The thermal optimum in the North Atlantic

129

independent foraminiferal-based proxies indicates that the general pattern of the

foraminifer-based temperature series is robust. Likewise, there is no a priori reason

to distrust the ﬁdelity of the diatom- and alkenone-based reconstructions. They

are consistent with summer season orbital forcing and consistent with other

surface ocean data from the region sampled from different cores, from ice-core

data, ice-core borehole data, and terrestrial records.

We must assume that both categories (i.e. with and without a pronounced early

to mid-Holocene thermal optimum) of temperature histories are valid and reﬂect

different aspects of the Holocene thermal evolution of the high-latitude ocean.

To understand why the difference has developed, the seasonality of the forcing,

the habitat of the biological proxy indicators, and the seasonality of the vertical

structure of the upper layers of the high-latitude ocean must be considered. The

main character of the orbital forcing of the early versus late Holocene is a strong

high-latitude early to mid-Holocene positive thermal anomaly during the summer

season, due to the combined effects of tilt and precession (see Cruciﬁx this volume,

Figure 4.4). A signiﬁcant, yet smaller, negative anomaly occurred in the winter

season, but is more conﬁned to low and mid-latitudes.

The result is an enhanced seasonality in the forcing and in the expected ocean

response as compared with the late Holocene. The discrepancy between the two

categories of temperature series might be explained as a response to the seasonality

of the forcing. Yet it remains to be argued why the foraminiferal-based records

should respond to the winter-time or annual mean forcing and not the summer-

time forcing, when we know that the primary foraminifer production season is in

the spring and summer. It is, however, well known that different species of

foraminifers have their main habitat at different depths in the ocean surface layer.

In contrast to diatoms and alkenone producing algae, which photosynthesize and

live in the euphotic zone of the upper 50 m of the ocean, planktonic foraminifers

are zooplankton and are often found at the depths of the main food sources near

the thermocline at the base of the surface layer. Risebrobakken et al. (2003) showed

that both the left- and right-coiling varieties of Neogloboquadrina pachyderma

share this pattern, i.e. showing a lack of evidence for the Holocene thermal opti-

mum, yet there is a difference between their O-isotope values, with the left-coiling

variety indicating colder temperatures than the right-coiling variety.

The seasonal forcing of surface layer temperatures provides strong contrasts

between the upper 50 m and deeper layers. Temperature data from Ocean Weather

Station M in the Norwegian Sea (Figure 5.6) shows that the summer solar insola-

tion warms and stratiﬁes the surface layer, imposing a strong heating in the upper

50 m, whereas the temperature of the thermocline is unrelated to the summer

season, and instead reﬂects the winter-time ventilation of the whole surface layer

when the summer season stratiﬁcation breaks down. Thus at 100 m, summer tem-

peratures are set by the winter-time situation and reﬂect this or the annual mean

temperatures. As can been seen in the hydrographic data (Figure 5.6), a tempera-

ture proxy derived from organisms that thrive below approximately 50 m in

these waters will not register any spring or summer-time warming, even though

they have their largest standing stock during the summer season. Enhanced solar

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