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Journal of Atmospheric and Solar-Terrestrial Physics 67 (2005) 55–69 The climate during the Maunder Minimum: a simulation with the Freie Universita¨t Berlin Climate Middle Atmosphere Model (FUB-CMAM) Ulrike Langematz , Antje Claussnitzer, Katja Matthes, Markus Kunze Institut fu¨r Meteorologie, Freie Universita¨t Berlin, Carl-Heinrich-Becker-Weg 6-10, 12165 Berlin, Germany Available online 2 September 2004 Abstract A model simulation of the climate during Maunder Minimum (MM) (1645–1715) was performed using the Freie Universita¨t Berlin Climate Middle Atmosphere Model (FUB-CMAM). A multi-year equilibrium integration with prescribed solar insolation, atmospheric composition and sea surface temperatures (SSTs) for MM conditions was compared with a present-day (PD) simulation. We find that during MM the stratosphere was significantly warmer (up to 3 K) than during PD, and dynamically more disturbed in winter. The warming is due to the dominant effect of the lower atmospheric CO 2 concentration during MM, which leads to a reduced emission of long-wave radiation, and compensates the cooling due to the reduced solar irradiance. The troposphere was about 1–1.5 K cooler in the annual mean during MM. The global mean surface air temperature decreased by 0.86 K. Northern hemisphere winters were on average characterized by cooler and drier weather over the northern parts of the continents, with an increase in precipitation in the southern parts. These climate anomalies are shown to be related to a shift in the North Atlantic Oscillation (NAO) towards a predominantly low phase during MM. The simulated climate anomalies are in very good agreement with reconstructions from proxy-data. Changes in the dynamical coupling between the troposphere and stratosphere were found in the MM simulation, indicating the importance of the stratosphere for climate change. r 2004 Elsevier Ltd. All rights reserved. Keywords: Past climate change; Solar activity; General circulation model; Stratospheric–tropospheric coupling 1. Introduction The potential impact of long-term solar irradiance changes is of particular interest for the interpretation of the warming trend in the 20th century. A pronounced, long-lasting anomaly in solar irradiance was the so- called Maunder Minimum (hereafter referred to as MM) during the Little Ice Age. In this paper, we study the climate of the MM by imposing solar irradiance and atmospheric composition changes for the MM on the Freie Universita¨t Berlin Climate Middle Atmosphere Model (FUB-CMAM). This study is the first of a series in which we intend to isolate the relative contributions of solar forcing, greenhouse gas and volcanic forcing on the reconstructed climate change during MM. The MM was a period characterized by a severe reduction of sunspot activity between about 1645 and 1715 (Eddy, 1976; Ribes and Nesme-Ribes, 1993). The 11-year Schwabe cycle of solar activity was still existent, however greatly suppressed (Beer et al., 1985; Usoskin et al., 2000), and reconstructions of solar activity using ARTICLE IN PRESS www.elsevier.com/locate/jastp 1364-6826/$ - see front matter r 2004 Elsevier Ltd. All rights reserved. doi:10.1016/j.jastp.2004.07.017 Corresponding author. Tel.: +30-838-711-65; fax: +30- 838-711-28. E-mail address: [email protected] (U. Langematz).

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ARTICLE IN PRESS

1364-6826/$ - se

doi:10.1016/j.ja

�Correspond838-711-28.

E-mail addr

(U. Langematz

Journal of Atmospheric and Solar-Terrestrial Physics 67 (2005) 55–69

www.elsevier.com/locate/jastp

The climate during the Maunder Minimum: a simulation withthe Freie Universitat Berlin Climate Middle Atmosphere

Model (FUB-CMAM)

Ulrike Langematz�, Antje Claussnitzer, Katja Matthes, Markus Kunze

Institut fur Meteorologie, Freie Universitat Berlin, Carl-Heinrich-Becker-Weg 6-10, 12165 Berlin, Germany

Available online 2 September 2004

Abstract

A model simulation of the climate during Maunder Minimum (MM) (1645–1715) was performed using the Freie

Universitat Berlin Climate Middle Atmosphere Model (FUB-CMAM). A multi-year equilibrium integration with

prescribed solar insolation, atmospheric composition and sea surface temperatures (SSTs) for MM conditions was

compared with a present-day (PD) simulation. We find that during MM the stratosphere was significantly warmer (up

to 3K) than during PD, and dynamically more disturbed in winter. The warming is due to the dominant effect of the

lower atmospheric CO2 concentration during MM, which leads to a reduced emission of long-wave radiation, and

compensates the cooling due to the reduced solar irradiance. The troposphere was about 1–1.5K cooler in the annual

mean during MM. The global mean surface air temperature decreased by 0.86K. Northern hemisphere winters were on

average characterized by cooler and drier weather over the northern parts of the continents, with an increase in

precipitation in the southern parts. These climate anomalies are shown to be related to a shift in the North Atlantic

Oscillation (NAO) towards a predominantly low phase during MM. The simulated climate anomalies are in very good

agreement with reconstructions from proxy-data. Changes in the dynamical coupling between the troposphere and

stratosphere were found in the MM simulation, indicating the importance of the stratosphere for climate change.

r 2004 Elsevier Ltd. All rights reserved.

Keywords: Past climate change; Solar activity; General circulation model; Stratospheric–tropospheric coupling

1. Introduction

The potential impact of long-term solar irradiance

changes is of particular interest for the interpretation of

the warming trend in the 20th century. A pronounced,

long-lasting anomaly in solar irradiance was the so-

called Maunder Minimum (hereafter referred to as MM)

during the Little Ice Age. In this paper, we study the

e front matter r 2004 Elsevier Ltd. All rights reserve

stp.2004.07.017

ing author. Tel.: +30-838-711-65; fax: +30-

ess: [email protected]

).

climate of the MM by imposing solar irradiance and

atmospheric composition changes for the MM on the

Freie Universitat Berlin Climate Middle Atmosphere

Model (FUB-CMAM). This study is the first of a series

in which we intend to isolate the relative contributions

of solar forcing, greenhouse gas and volcanic forcing on

the reconstructed climate change during MM.

The MM was a period characterized by a severe

reduction of sunspot activity between about 1645 and

1715 (Eddy, 1976; Ribes and Nesme-Ribes, 1993). The

11-year Schwabe cycle of solar activity was still existent,

however greatly suppressed (Beer et al., 1985; Usoskin et

al., 2000), and reconstructions of solar activity using

d.

ARTICLE IN PRESSU. Langematz et al. / Journal of Atmospheric and Solar-Terrestrial Physics 67 (2005) 55–6956

historical proxies suggest a decrease of total solar

irradiance during the MM ranging from 0.05% to

0.5% (Hoyt and Schatten, 1993; Lean et al., 1995; Lean,

2000) compared to the present day (hereafter referred to

as PD).

This minimum in solar activity coincided with the

coldest period of the Little Ice Age. Historical recon-

structions of Northern Hemisphere (NH) surface air

temperatures from different proxies such as tree rings

and ice cores show that the strongest cooling of the past

millenium occurred in the 17th century. Negative

temperature anomalies reached 0.6K in the annual

mean compared to the reference period 1901–1950

(Overpeck et al., 1997; Jones et al., 1998; Mann et al.,

1999; Briffa, 2000; Briffa et al., 2001; Crowley and

Lowery, 2000; Esper et al., 2002). The winters and

springs in continental western Europe were cold and dry

with temperatures locally reduced by 1–1.5K (e.g.,

Manley, 1974; Glaser and Hagedorn, 1991; Pfister,

1992), while summers were only slightly colder than in

the first half of the 20th century (Manley, 1974; Pfister,

1995). Using multi-proxy-data and documentary records

Wanner et al. (1995) were able to reconstruct seasonal

synoptic sea level pressure maps over Europe for the late

MM (1675–1704). They showed that during that period

blockings over north-west Europe were quite frequent in

winter, leading to southward outbreaks of cold,

continental air. The pressure dipole with high pressure

over the North Atlantic/western Europe and low

pressure over south-west Europe is equivalent to a low

North Atlantic Oscillation (NAO) index. Independent

reconstructions of the NAO index from different proxy-

data (Appenzeller et al., 1998; Glueck and Stockton,

2001; Luterbacher et al., 2002) indicate indeed that, in

spite of its large interannual variability, the NAO index

was predominantly negative during MM. In contrast,

PD winters more often feature a high NAO index with a

pronounced Icelandic low and Azores high, and a more

zonal circulation. The spring season during MM was

dominated by troughs over western Europe and

meridional flow with strong cold air outbreaks from

the North Sea towards southern Europe. European

summers could feature strong hailstorms and frosts

(Pfister, 1994), whereas proxy-data revealed a prolonged

dry period over Canada (George and Nielsen, 2002) and

the northern United States (Cook et al., 1999).

LaMarche (1974) inferred a cold and dry summer

climate in south-western United States during MM.

While the coincidence between the reconstructed solar

and climate anomalies during the MM suggests some

connection between both, the proof of a causal relation-

ship requests model simulations in which the relevant

processes are incorporated. A number of modeling

studies investigated the impact of variations in solar

activity. First estimates by Kelly and Wigley (1992) and

Schlesinger and Ramankutty (1992) indicated a partial

solar climatic impact on the observed surface tempera-

ture increase since the 18th century. Rind and Overpeck

(1993) simulated the equilibrium response to a 0.25%

solar irradiance reduction (Lean et al., 1992) using the

tropospheric, coarse-grid version of the Goddard

Institute for Space Studies (GISS)-GCM; they found a

global annual mean cooling of 0.45K. A similar

response was obtained when forcing a global coupled

ocean–atmosphere model with a time-dependent varia-

tion in solar irradiance of 0.35% (Hoyt and Schatten,

1993): a cooling of the global mean surface air

temperature of about 0.5K (Cubasch et al., 1997). More

recently, Rind et al. (1999) studied the time-dependent

climate response to a 0.25% decrease in solar irradiance

during MM, again with the tropospheric GISS-GCM,

however, this time coupled to a mixed-layer (slab) ocean

module. Their results showed a cooling of 0.45K due to

the reduced solar forcing.

Given the uncertainties in the prescribed solar forcing

and the differences in the GCMs and performed

simulations, there is surprisingly good agreement in

the annual and global mean temperature responses.

However, the regional response patterns differ between

the models. So, the increase of surface air temperature

over the North Atlantic produced by several GCMs is

no longer existent in the recent GCM study of Fischer-

Bruns et al. (2002), who used an updated version of the

coupled ocean–atmosphere GCM of Cubasch et al.

(1997). Possible explanations for the simulated discre-

pancies between models are the representation of the

middle atmosphere which is needed to account for

the impact of solar variability on ozone, as well as the

consideration of the spectral variation of solar varia-

bility. Lean (2000) estimated a decrease in the short-

wave (SW) UV radiation of several percent during MM,

which affects the radiative balance of the stratosphere as

well as the ozone concentration (Haigh, 1994; Shindell et

al., 1999). Matthes et al. (2003) showed that GCMs

considering spectral variations between different phases

of the 11-year solar cycle show a clear direct solar signal

in the temperature and circulation of the upper strato-

sphere. The signal may be transferred into the lower

stratosphere and troposphere (e.g., Kodera and Kuroda,

2002; Matthes et al., 2004) and have an impact on

climate. To account for these effects, GCMs that resolve

the atmosphere from the Earth’s surface to the upper

mesosphere with high-resolution SW radiation schemes

should be used. Another factor influencing the models’

response to solar forcing is the degree of their

ocean–atmosphere coupling. Different studies indicated

a close relationship between solar irradiance changes

and the oceans, as e.g., Bond et al. (2001), who found on

the time scale of millenia that cooler surface water and

increased southward transport of drift-ice due to

enhanced winds from the north are correlated

with reduced solar forcing. To date, fully coupled

ARTICLE IN PRESSU. Langematz et al. / Journal of Atmospheric and Solar-Terrestrial Physics 67 (2005) 55–69 57

atmosphere–ocean GCMs are restricted to the tropo-

sphere and lower stratosphere, as e.g., Fischer-Bruns et

al. (2002). Limited computer resources did not allow to

develop fully coupled middle atmosphere (MA)–ocean

models. Instead, MA GCMs use either simplified mixed-

layer ocean modules or prescribed sea surface tempera-

tures (SST), as in this study.

Shindell et al. (2001) were the first to use spectrally

discriminated irradiances in a middle atmosphere ver-

sion of the GISS-GCM. They compared MM conditions

(around 1680) with pre-industrial conditions 100 years

later. The increase in solar irradiance from 1680 to 1780

in their model led to a global annual mean warming of

0.34K. Regional climate anomalies could be up to five

times greater and resembled a shift towards the low AO/

NAO index, in good agreement with reconstructions

from proxy-data (e.g., Appenzeller et al., 1998; Luter-

bacher et al., 2002).

The goal of this study is to find out whether the

reconstructed climate anomalies during MM can be

reproduced in a model simulation when considering the

estimated changes of external solar forcing, atmospheric

composition and related ocean anomalies between MM

and PD. In contrast to previous model studies, we used a

troposphere–stratosphere–mesosphere GCM which al-

lows to study the coupling processes between the

troposphere and the middle atmosphere from the

Earth’s surface up to the mesopause. A further

improvement refers to a high-resolution SW radiation

code which uses the most up-to-date estimates of

spectral solar variations. The model and the experi-

mental design are introduced in Section 2. The strato-

spheric signal as the primary response to the external

forcing in the middle atmosphere will be analyzed first in

Section 3. In Section 4 we will present the simulated

near-surface climate signal and compare it with recon-

structions from proxy-data. The coupling between the

stratosphere and the troposphere will be addressed in

Section 5. Conclusions and a discussion will follow in

Section 6.

2. Model and experiments

2.1. Model description

The model used for this study is the FUB-CMAM. It

is run at a T21 spectral resolution corresponding to a

horizontal resolution of 5:625� 5:625� in gridpoint

space. The model has 34 levels with the lid at

0.0068 hPa ð� 83kmÞ and a vertical resolution in the

middle atmosphere of 3.5 km. It includes physical

parametrizations of the hydrological cycle and vertical

diffusion in the troposphere, as well as a Rayleigh

friction in the upper mesosphere to account for the

effects of breaking gravity waves. Further details are

given in Pawson et al. (1998) and Langematz (2000).

The FUB-CMAM includes a state-of-the-art radia-

tion scheme for absorption and emission due to carbon

dioxide (CO2), ozone (O3) and water vapor (H2O)

(Morcrette, 1991). In the basic model version, absorp-

tion of SW solar radiation in the ultraviolet (UV) and

visible (VIS) spectrum by O3 and oxygen (O2) at heights

above 70 hPa was calculated using the parametrizations

of Shine and Rickaby (1989) and Strobel (1978). To

account for spectral changes in solar insolation between

the MM period and PD, we used here an improved SW

radiation code in which the spectral resolution between

206.2 and 852.5 nm was increased from 8 bands (of the

Shine and Rickaby scheme) to 44 bands, (Matthes et al.,

2004) based on high-resolution line intensities and

absorption coefficients from WMO (1986). A climato-

logical zonal and monthly mean O3 distribution as well

as a global mean CO2 mixing ratio are prescribed. Note

that the radiative effect of methane changes during MM

had to be neglected in our simulation due to the nature

of the radiative code.

As the model version used here does not include a

sophisticated gravity wave scheme in the mesosphere, it

does not simulate the Quasi Biennial Oscillation (QBO)

of the tropical winds in the stratosphere.

2.2. Experimental design

To investigate the changes between the MM period

and PD the statistical results of two equilibrium states

simulated by the FUB-CMAM are compared, one

representing MM conditions, the other representing

PD conditions.

MM conditions are considered in the model by

changes in solar insolation, chemical composition and

lower boundary forcing. To account for the decrease in

insolation during MM, high-resolution spectral solar

flux changes estimated by Lean (2000) were adapted to

the 44 spectral bands of the FUB-CMAM SW radiation

scheme. This leads to a decrease from about 1% between

300 and 400 nm to 6% between 250 and 210 nm.

Additionally, the total solar insolation was reduced by

0.2% (Lean, 2000).

Changes in the ozone concentration between MM and

PD were implemented by considering two impacts: due

to the decrease in insolation, photochemical ozone

production was reduced during MM. These solar

induced ozone changes were precalculated in a two-

dimensional chemical transport model (Haigh, pers.

comm.). In contrast, stratospheric ozone levels were

higher during MM than PD because the concentrations

of chlorine, hydrogen and nitrogen oxides were lower in

the pre-industrial atmosphere leading to a reduced

catalytic ozone destruction. This effect was simulated

in a two-dimensional chemical-transport model by

ARTICLE IN PRESSU. Langematz et al. / Journal of Atmospheric and Solar-Terrestrial Physics 67 (2005) 55–6958

Wuebbels et al. (1998) to cause an ozone increase during

MM in the upper stratosphere of more than 10% as

compared to the present, with stronger local enchance-

ments due to transport. For our study, an annual mean,

latitudinally homogeneous percentage ozone increase

was estimated from Wuebbels et al. (1998). The increase

extends from the middle stratosphere to the lower

mesosphere and reaches its maximum of about 10%

around 45 km. Both ozone changes (solar decrease and

pre-industrial increase) were imposed on a 1980s back-

ground ozone distribution, constructed from the CIRA

climatology and SBUV satellite measurements (Kubitz,

pers. comm.). This basic ozone field was chosen to

exclude the Antarctic ozone hole in the lower strato-

sphere (occurring since the late 1970s) from the

‘unpolluted’ atmosphere during MM. Fig. 1 shows the

resulting annual mean ozone changes when both solar

minimum and pre-industrial conditions are considered.

The ozone concentration during the MM was higher in

the upper stratosphere by up to 8% around the

stratopause (pre-industrial effect) and lower in the

middle and lower stratosphere (at heights below

35 km) reaching �5% at middle and high latitudes

(solar effect). The maximum solar-induced ozone

decrease in the upper stratosphere of about 6% is more

than compensated by the pre-industrial increase. The

ozone profile changes sign at approximately the height

of the maximum in ozone mixing ratio so that during

MM the maximum ozone concentration was located

higher in the stratosphere than in PD. In addition, the

CO2 amount was decreased in the MM simulation to

pre-industrial levels (280 ppmv). The impact of volcanic

forcing was neglected in this study.

To allow for a response of the oceans to MM

conditions, the prescribed SST distribution was modified

by changes calculated with a coupled ocean–atmosphere

model of the Deutsches Klimarechenzentrum (DKRZ)

Fig. 1. Annually averaged zonal mean percentage change from

PD to MM of the ozone volume mixing ratio (contour interval

2%). Negative values are shaded.

(Fischer-Bruns et al., 2002). SST anomalies, from their

transient simulation which were averaged for the period

1645–1715, were used. They were negative over large

areas except for southern middle-to-high latitudes.

Rather strong cooling occurred in the North Atlantic

ocean in winter, with a maximum of up to 4K south of

Greenland and in the north-western Pacific. In contrast,

the cooling of the polar oceans was weaker during

summer, with a slight SST increase south of Greenland.

The tropical oceans were cooler by about 0.5–0.75K.

The PD simulation is representative for current

atmospheric conditions. For comparison with the MM

run, we use the mean response of two equilibrium PD

simulations, one using solar insolation for maximum

conditions of the 11-year solar cycle, and the other using

minimum conditions. The ozone distribution of the PD

runs was constructed from satellite and ozonesonde

measurements and represents the 1990s, including the

Antarctic ‘ozone hole’ in spring (updated from Fortuin

and Langematz (1994)). Climatological mean SSTs

based on the 1979–1994 AMIP (Atmospheric Modelling

Intercomparison Project) data set were prescribed. The

CO2 concentration was 330 ppmv. The simulation of the

11-year solar signal in those runs were discussed in detail

in Matthes et al. (2003).

3. Stratospheric response

Fig. 2 shows the zonal and annual mean thermal and

dynamical response of the model atmosphere to the

implied forcing. Three areas with significant temperature

changes emerge (Fig. 2a): (1) The global upper strato-

sphere and lower mesosphere were warmer during MM

reaching a maximum heating of 3K at the stratopause.

(2) The lower polar stratosphere was warmer during

MM, in particular, in the Southern Hemisphere where

temperatures were higher by about 2.5K compared to

PD. (3) The global troposphere featured lower zonal

mean temperatures up to 1K during MM. The strongest

cooling of 1.5K occurs in the tropical and subtropical

upper troposphere and extends into the lower strato-

sphere and leads to a rising of the tropopause.

The described temperature changes are connected to

circulation changes (Fig. 2b). In the annual mean, the

SH mesospheric jet is significantly weakened by about

3m/s. The upper troposhere/lower stratosphere wester-

lies are reduced on both hemispheres between 2m/s in

the NH and 4.5m/s in the SH. Significant wind changes

extend down to the Earth’s surface where slightly

stronger zonal mean zonal winds in the tropical and

subtropical latitudes are neighbored by weaker winds in

mid-latitudes of both hemispheres.

As the implied forcing for MM, i.e. changes in solar

irradiance, ozone and carbon dioxide, directly affects the

stratospheric radiation budget, we examined the change

ARTICLE IN PRESS

(a)

(b)

Fig. 2. Annually averaged zonal mean changes from PD to

MM of: (a) temperature (contour interval 0.5K) and (b) zonal

wind (contour interval 0.5m/s). Dark (light) shaded areas

denote regions where the changes are significant at the 99%

(95%) level.

Fig. 3. Annually averaged zonal mean change in SW heating

rate from PD to MM (contour interval 0.1 K/day). Dark (light)

shaded areas denote regions where the changes are significant at

the 99% (95%) level.

U. Langematz et al. / Journal of Atmospheric and Solar-Terrestrial Physics 67 (2005) 55–69 59

in the annual mean heating rates between the MM and

PD experiments. The SW heating rate differences,

shown in Fig. 3, display a layered structure throughout

the atmosphere with negative SW heating rates dom-

inating. This is mainly the result of the decrease in solar

irradiance during MM. In a separate sensitivity study, in

which only solar irradiance was reduced, we found a

decrease of the tropical solar heating rate of around

60 km of � 0:09K=day from maximum to minimum

insolation during the 11-year solar cycle. The SW

heating rate decrease during MM of 0.2K/day is twice

as large in the same height region (Fig. 3). Given the

almost twice as large decrease of solar insolation during

MM, this indicates a direct effect of the missing

insolation. Below 60 km, the changed ozone concentra-

tion additionally modifies the radiation budget. The

upper stratosphere between 38 and 60 km featured more

ozone during MM, while less ozone was present in the

lower stratosphere (Fig. 1). However, the anticipated

increase in SW heating rates in the upper stratosphere is

more than compensated by the direct irradiance effect,

except for the stratopause region, where the ozone effect

is strongest and reaches an SW heating of 0.1K/day. In

the middle and lower stratosphere below 35 km both

effects act in the same direction, namely to cool the

stratosphere.

When comparing the SW heating rate changes in

Fig. 3 with the zonal mean temperature response

(Fig. 2a), it becomes very clear, that the warming of

the upper stratosphere cannot be due to the solar

irradiance and ozone changes during MM. The stron-

gest effect, responsible for the warming of the strato-

sphere, is the lower pre-industrial CO2 amount.

Langematz et al. (2003), by using the same GCM to

study the impact of anthropogenic CO2 changes,

calculated a temperature decrease around the strato-

pause of � 1K for a CO2 increase of 15 ppmv between

1980 and 2000. Correspondingly, the implied CO2

decrease between PD and MM of 50 ppmv should lead

to a warming of the same layer by about 3.3K.

Although the long-wave (LW) cooling is not only a

function of the gas concentration but also highly

dependent on the thermal structure, the good corre-

spondence between this rough estimate and the actual

warming of 3K (Fig. 2a) strongly suggests that the

reduction in LW emission due to the lower CO2 amount

during MM is responsible for the warming of the upper

stratosphere. The warming region between 10 and 20 km

over Antarctica is caused by the higher MM ozone

concentration (non-existent heterogeneous ozone de-

struction by CFCs during MM, which is reflected in the

prescribed ozone distributions, not shown), while the

cooling of the lower stratosphere and the troposphere

are a composite effect of the insolation, ozone and CO2

(smaller greenhouse effect) changes. There are also

indications that the dynamical variability in the

stratosphere was enhanced during MM. Both model

ARTICLE IN PRESSU. Langematz et al. / Journal of Atmospheric and Solar-Terrestrial Physics 67 (2005) 55–6960

simulations, for MM and for PD, display the observed

interannual variability during the NH winter season:

almost each winter is characterized by sudden strato-

spheric warmings leading to a weakening of the strato-

spheric vortex and of the polar night jet. However, in the

MM simulation a shift towards an earlier occurrence of

strong minor warmings exists, with two exceptionally

strong minor warmings in early winter (December,

January) during the 15 years of simulation. However,

more simulation years would be needed to test the

significance of this result.

4. Climate response

In this section we study the long-term mean response

of near-surface variables to the implied MM forcing and

compare the results with available reconstructions from

proxy-data, as well as with other model estimates.

4.1. Annual mean changes

In Fig. 4 we show the NH annual mean surface air

temperature change between MM and PD. The North-

ern Hemisphere is dominated by a cooling of �1 to

�2K, except for three localized warming areas: a weak

warming stretching along the west coast of the North

American continent, a second area centered over New-

foundland reaching 2K, and a third one over Kamchat-

ka with values of up to 1K. Except for the North

American warming region and some small local areas

over the South Pacific, the temperature signal is

statistically significant everywhere at the 99% confi-

dence level, calculated with a Student’s t-test. It thus

exceeds significantly the internal variability of the PD

simulation and can be regarded as a response to the

prescribed changes for the MM.

longitude

latit

ude

Fig. 4. Annual mean change from PD to MM of temperature

at the 1000hPa pressure level (contour interval 1K; the �0:5Kcontour lines are shown as well). Negative values are shaded.

Averaged over the NH, the FUB-CMAM calculates a

surface cooling during MM of 0.86K compared to PD,

which is in very good agreement with the estimate of a

0.6–1K cooling from proxy-data, as summarized by

Palmer (2002). Note, however, that the range of

reconstructed cooling varies from 0.25K below the

1902–1980 average (Mann et al., 1998, 1999) to 0.8K

below the 1961–1990 average (Briffa et al., 2001) to 1K

compared to the 20th century (Esper et al., 2002). The

simulated cooling is on the ‘cool’ side of the reconstruc-

tions which can be explained by the experimental setup;

while the reconstructed cooling is generally referred to

an average period of the 20th century, the model’s PD

baseline is an equilibrium 1990 climatology. For north-

western Europe, the simulated cooling is about 1K,

similar to the 1K cooling derived from proxy-data

(Pfister, 1992). The North American continent has the

strongest cooling of 3.5K over north-eastern Canada,

qualitatively consistent with reconstructions (e.g.,

LaMarche, 1974).

The SH is dominated by a cooling in subtropical and

middle latitudes reaching 1 to 1.5K over the continents.

At higher southern latitudes the influence of the higher

SSTs leads to a weak regional warming of the surface air

temperature during MM.

The magnitude of the annual mean cooling over the

tropical and subtropical oceans and in the SH mid- to

high latitudes suggests a large influence of the imposed

SST changes. This is confirmed by the good agreement

in those regions with the results of Fischer-Bruns et al.

(2002), who provided the MM SST anomalies for our

study. However, the NH regional response patterns of

surface air temperature are different from Fischer-Bruns

et al. (2002), with the positive anomalies of our

simulation missing in Fischer-Bruns et al. (2002). This

suggests that the SST anomalies are not the only factor

for the tropospheric climate change but, as suggested by

Shindell et al. (2001), the coupling with the stratosphere

in the FUB-CMAM provides additional impact. How-

ever, a definite conclusion of the relative contributions

can only be drawn from additional GCM runs with

separate forcings.

4.2. Seasonal mean changes

As shown in Fig. 5, the annual mean changes in

surface air temperature (Fig. 4) are a composite of the

seasonal signals in NH winter (December/January/

February, DJF) and summer (June/July/August, JJA).

Although the NH average temperature change is almost

identical for both seasons (�0:86K in DJF; �0:84K in

JJA), more extreme changes occur in winter and

determine the regional patterns of the annual mean

response. In DJF (Fig. 5a), positive temperature

anomalies over Newfoundland and Kamchatka reach

4K, while the Canadian Arctic cools up to 6K. In

ARTICLE IN PRESS

(a) (b)

Fig. 5. Seasonal mean change from PD to MM of the 1000 hPa temperature for: (a) NH winter (DJF) and (b) NH summer (JJA)

between 20�N and 90�N (contour interval 1K). Dark (light) shaded areas denote regions where the changes are significant at the 99%

(95%) level.

U. Langematz et al. / Journal of Atmospheric and Solar-Terrestrial Physics 67 (2005) 55–69 61

contrast, NH summer is characterized by a more

uniform cooling of the land masses between 1 and 2K

(Fig. 5b). Only the warming area over Newfoundland

persists into summer reaching a positive anomaly of 1K.

However, the enhanced amplitude and the pattern of the

regional temperature signal in winter strongly suggest a

dynamical feedback.

Therefore, we focus now on the NH winter season.

Fig. 6 displays the DJF anomalies of mean sea level

pressure (Fig. 6a), the horizontal wind vector (Fig. 6b),

and the zonal (Fig. 6c) and meridional (Fig. 6d) wind

components at 1000 hPa. Four centers of strong and

significant sea level pressure anomalies appear (Fig. 6a),

with positive anomalies over the North Atlantic

(+3.5 hPa) and east Siberia (+2.5 hPa) and negative

anomalies over the east Atlantic and the Mediterranean

ð�2:5hPaÞ and the north-eastern Pacific ð�3hPaÞ. The

Atlantic pressure anomalies correspond to a reduction in

the intensity of the Icelandic low and the Azores high

during MM. The North Pacific pressure anomalies are

anticorrelated with those of the North Atlantic, while

they are correlated with those over the Mediterranean.

This sea-saw is a well-known phenomenon (Van Loon

and Rogers, 1978) and indicates a shift towards a

different circulation mode in MM compared to PD. The

reduced meridional pressure gradient over the North

Atlantic is associated with circulation changes in mid-

latitudes. The westerly flow is reduced by about 1.5m/s

in northern winter (Fig. 6c), while the meridional

component shows a southward anomaly of more than

1m/s (Fig. 6d). Even stronger and statistically significant

cyclonic anomalies of the horizontal wind occur over the

Pacific due to a pronounced decrease in mean sea level

pressure (Fig. 6a). The composite horizontal wind

anomalies clearly display the centers of flow changes

over the northern oceans.

Together with the circulation, precipitation is changed

in NH winter during MM, as shown in Fig. 7 for the

composite of large-scale and convective precipitation,

and snowfall. Large areas over northern/central Europe,

the North Atlantic and North America featured drier

winters in MM, with locally statistically significant

decreases in large-scale precipitation, e.g., over Scandi-

navia and the North Atlantic (not shown). Changes of

the opposite sign occurred over the North Atlantic south

of 40�N, the Mediterranean, as well as over the North

Pacific and the western United States. Here, winters

were more wet, due to combined, locally significant

increases in large-scale and convective precipitation. A

similar increase in precipitation was derived in a

reconstruction of rainfall anomalies for southern Spain

(Rodrigo et al., 2000). Luterbacher and Xoplaki (2003)

found that the wettest winter of a 500-year time series

occurred in 1684, during Late MM, when Portugal, Italy

and the western Balkans obtained 30–90mm more

precipitation than usual. In the GCM, we obtain about

30mm/month more rainfall in winter over the eastern

Mediterranean during MM. Further, a pronounced

increase in convective precipitation was simulated over

Newfoundland.

Changes in the intensity of the Icelandic low and the

high pressure over the east Atlantic and south-western

Europe explain to a large extent the variability of NH

winter (Barnston and Livezey, 1987). One quantity to

describe the phase of this pressure dipole is the NAO

index i, which is defined as normalized pressure

difference between two geographical locations in the

centers of the pressure systems (Hurrell, 1995). In Fig. 8

the time series of the NAO index for each winter

(averaged over DJF) of the MM (Fig. 8a) and PD (Fig.

8b) simulations are shown. Both simulations display

considerable interannual variability, with altering peri-

ARTICLE IN PRESS

(a) (b)

(c) (d)

Fig. 6. Change from PD to MM of: (a) mean sea level pressure (contour interval 0.5 hPa), (b) the horizontal wind vector, (c) the zonal

wind component (contour interval 0.4m/s), and (d) the meridional wind component (contour interval 0.4m/s) at 1000 hPa for NH

winter (DJF) between 20�N and 90�N. Dark (light) shaded areas denote regions where the changes are significant at the 99% (95%)

level.

U. Langematz et al. / Journal of Atmospheric and Solar-Terrestrial Physics 67 (2005) 55–6962

ods of either positive or negative NAO index. In the

MM simulation, however, more years with negative

NAO index occur than in the PD simulation. The

averaged NAO index is negative for MM ði ¼ �0:66Þ,while it is positive for PD ði ¼ 0:50Þ. A lower NAO index

is equivalent to a weaker pressure dipole, weaker

westerly flow towards Europe with reduced storm

activity and reduced advection of warm and moist air

from the Atlantic. Note that due to the wide spread of

NAO index values and the limited number of model

years, the NAO index changes are not statistically

significant. However, significant changes of mean sea

level pressure over the North Atlantic and south-western

Europe are obvious (Fig. 6a), upon which the NAO

index calculations are based. The simulated shift to

weaker NAO phases in MM agrees very well with

reconstructions from different types of proxy-data.

Appenzeller et al. (1998), Glueck and Stockton (2001)

and Luterbacher et al. (2002) showed that although the

NAO index was characterized by a high interannual

variability, it was predominantly in its low phase

during MM until about 1700. Based on paleoclimate

records, Bond (pers. comm.) also suggests a

recurring shift towards a negative NAO-like

pattern during reductions in solar activity for the MM

as well as for the past 12,000 years. Our model

result is also in good agreement with the model

study of Shindell et al. (2001), who found a similar shift

of the Arctic Oscillation (including the NAO) towards

lower states for the MM compared to 1780 pre-

industrial conditions. In contrast, Fischer-Bruns et al.

(2002) do not show a persistent negative NAO index

during MM, but instead an early transition towards the

positive NAO phase during the Late MM (LMM).

ARTICLE IN PRESSU. Langematz et al. / Journal of Atmospheric and Solar-Terrestrial Physics 67 (2005) 55–69 63

However, they analyze LMM in a transient model

simulation, while we compare a mean MM equilibrium

state (including the early MM) with equilibrium PD

conditions.

The simulated surface air temperature anomalies in

MM during NH winter (Fig. 5a) with negative

anomalies over the North American and Eurasian

continents and large positive anomalies over Newfound-

land are consistent with the high correlation between

temperature and the NAO index revealed by Hurrell and

van Loon (1997).

Fig. 8. NAO Index calculated from monthly mean sea level pressu

simulation (left) and the PD simulation (right). Straight lines indicate

Fig. 7. Change from PD to MM of precipitation (large-scale

plus convective plus snowfall) (contour interval 3� 10�9 m=s;1� 10�9 m=s is equivalent to 2.6mm/month) for NH winter

(DJF) between 20�N and 90�N. Negative values are shaded.

In summary, the model displays a consistent change

of the NH winter climate during MM featuring a

predominant low phase of the NAO, connected to colder

and drier winters in western and central Europe and

moister winters in southern Europe than PD.

5. Stratospheric–tropospheric coupling

In Section 3 we showed that the stratosphere was

significantly warmer and more disturbed during MM

compared to PD. The clear climate response to the MM

forcing with a more negative NAO was discussed in

Section 4. It is well known from a great number of

theoretical and numerical studies that the troposphere

influences the stratosphere by upward propagating

planetary waves, which in winter determine the strength

and variability of the stratospheric polar vortex (see

Andrews et al., 1987 and references therein). Further-

more, more recent analyses of observational data

suggested the existence of a statistical coupling between

tropospheric and stratospheric variability modes (Perl-

witz and Graf, 1995), which seems to be directed from

the stratosphere to the troposphere (Baldwin and

Dunkerton, 1999). This relationship can be described

in terms of the northern and southern annular modes

(NAM and SAM), which are defined as the leading

modes of geopotential variability (Baldwin and Dun-

kerton, 2001). At the surface, the NAM is identical to

the Arctic Oscillation (AO), whose Atlantic part is the

NAO. In the stratosphere the NAM is related to the

strength of the polar vortex. Baldwin and Dunkerton

(2001) showed that on average a weak stratospheric

vortex (negative NAM) is followed by a negative NAM

(AO, NAO) situation in the troposphere.

re in NH winter (DJF) for the individual years of the MM

mean values for the simulation periods.

ARTICLE IN PRESS

(a) (b)

(c) (d)

Fig. 9. Change from PD to MM of: (a) zonal wind (contour interval 0.5m/s), (b) the divergence of the EP flux (contour interval 1m/s/

day; the 0.5m/s/day contour line is shown additionally), (c) the vertical component of the EP flux (contour interval 1� 105 kg=s2, the0:5� 105 kg=s2, 0:3� 105 kg=s2, and 0:1� 105 kg=s2 contour lines are shown additionally), and (d) the meridional component of the

EP flux (contour interval 2� 105 kg=s2) for NH winter (DJF) in the troposphere and lower/middle stratosphere. Dark (light) shaded

areas denote regions where the changes are significant at the 99% (95%) level.

U. Langematz et al. / Journal of Atmospheric and Solar-Terrestrial Physics 67 (2005) 55–6964

In the following we examine whether a relationship

can be identified between the tropospheric and strato-

spheric signals in our MM simulation. As shown in

Fig. 9a, NH winter in the MM simulation is character-

ized by a statistically significant decrease of the zonal

mean zonal winds in mid-latitudes throughout the whole

troposphere and lower/middle stratosphere. In the

middle and upper troposphere, this weakening of the

westerly jet is associated with an enhanced wave activity,

as shown by the convergence of the Eliassen–Palm (EP)

flux (Fig. 9b, negative differences), which is a measure of

the interaction between propagating planetary or

synoptic scale waves and the mean flow. Significant

changes extend into the stratosphere as well: the reduced

westerlies in the upper troposphere lead to an enhanced

vertical propagation of tropospheric waves into the

stratosphere in mid- and high latitudes, as shown by the

increase of the vertical component of the EP flux

(Fig. 9c, positive differences), which is proportional to

upward propagating wave energy from the troposphere.

In the lower stratosphere, the waves then propagate

further southward during MM, as shown by the

meridional component of the EP flux (Fig. 9d, negative

differences south of 50�), which describes wave

propagation in the north–south direction. The

waves then dissipate in the lower and middle strato-

sphere (Fig. 9b) and decelerate the zonal mean wind

(Fig. 9a).

In contrast, in the upper stratosphere, planetary wave

activity is significantly weaker at mid- and high latitudes

during MM (not shown). This result is in very good

agreement with the doubled CO2 simulation of Rind et

al. (1998). While they found an enhanced wave activity

and dynamical warming of the upper stratosphere, when

CO2 is increased, we obtain the opposite result of

reduced wave activity and dynamical cooling, when CO2

is reduced. We therefore assume that the upper

stratosphere thermal and dynamical anomalies during

MM are predominantly due to the changed CO2

concentration.

In Fig. 10 we show the NH geopotential height

changes between MM and PD for NH winter (DJF) for

different constant pressure levels from the surface to

1 hPa (48 km). While the 1000 hPa level is dominated by

regional NAO-type pressure changes, as discussed in

Section 4, the change patterns become more annular in

ARTICLE IN PRESS

Fig. 10. Geopotential height changes from PD to MM for NH winter (DJF) at 1000, 500, 100, 30, 10, and 1 hPa (contour intervals

0.5 dam at 1000 and 500 hPa, 1 dam at 100 hPa, 2 dam at 30 and 10 hPa and 4 dam at 1 hPa) between 20�N and 90�N. Dark (light)

shaded areas denote regions where the changes are significant at the 99% (95%) level.

U. Langematz et al. / Journal of Atmospheric and Solar-Terrestrial Physics 67 (2005) 55–69 65

the mid- and upper troposphere. In the stratosphere a

wavenumber 1 pattern evolves with height, indicating a

weakening of the polar vortex and a shift towards

Europe. Thus, the weaker and more disturbed strato-

spheric polar vortex in NH winter and the negative

NAO index at the surface confirm the coupling between

troposphere and stratosphere, inferred from observa-

tional analyses by Perlwitz and Graf (1995) and Baldwin

and Dunkerton (1999, 2001). A direct comparison

would need the calculation of daily NAM indices for

both experiments which will be the subject of future

work.

ARTICLE IN PRESSU. Langematz et al. / Journal of Atmospheric and Solar-Terrestrial Physics 67 (2005) 55–6966

6. Conclusions and discussion

We described here results of a multi-year equilibrium

simulation of the MM climate using a GCM with a

resolved stratosphere and mesosphere. MM conditions

were prescribed to the model by including a spectrally

resolved decrease in solar irradiance, modifying the

atmospheric composition according to the changed

insolation and pre-industrial conditions, and adapting

the fixed SSTs at the lower boundary. The climate

response was isolated by comparing the statistically

mean long-term differences between the MM simulation

and a control run for PD conditions.

The main results for the MM are:

(a)

The stratosphere was warmer, reaching the maxima

of 2.5–3K at the stratopause and the SH polar lower

stratosphere in the annual mean. The troposphere

was cooler, reaching a maximum cooling of 1.5K in

the tropical/subtropical upper troposphere. The

westerly jets were reduced in the upper SH meso-

sphere and in the upper troposphere/lower strato-

sphere of both hemispheres.

(b)

An analysis of the stratospheric radiation budget

revealed that the temperature changes in the middle

and upper stratosphere were dominated by the lower

CO2 content, while the decrease in solar irradiance

and the induced ozone changes played a minor role.

In contrast, the warmer lower stratosphere at SH

high latitudes during MM was due to the higher

MM ozone concentration (no ozone hole in MM).

(c)

The surface climate was cooler: � 0:86K in the NH

and 1–1.5K in northern and central Europe. The

simulated temperature anomalies are highly signifi-

cant, and are in very good agreement with recon-

structions from different sources of proxy-data.

(d)

Regional climate anomalies were enhanced during

NH winter. The NAO was shown to be predomi-

nantly in its low phase, being equivalent to a reduced

Icelandic low and Azores high, with reduced

westerly circulation and colder and drier winters in

northern/central Europe. In contrast, enhanced

precipitation occurred in southern Europe and over

the North Pacific. The simulated anomaly patterns

agree well with well-known observed correlation

patterns and confirm the shift in the tropospheric

variability mode. These results again are in very

good agreement with reconstructions.

(e)

The tropospheric temperature anomalies led to an

enhanced propagation of tropospheric waves into

the lower/middle stratosphere during winter, where

they dissipate and decelerate the zonal mean flow,

while in the upper stratosphere the impact of the

CO2 changes dominates the dynamical response.

The weaker and more disturbed polar vortex in the

stratosphere in NH winter and the negative NAO

index at the surface suggest the existence of the

observed coupling between the stratosphere and the

troposphere.

In summary, given the good agreement between the

reconstructed climate signal during MM and our model

simulation, we conclude that the climate anomalies

during MM were mainly caused by a combination of the

change in solar forcing (with resulting SST changes) and

the different atmospheric composition during the 17th

century (with more stratospheric ozone and less CO2).

In order to estimate the relative contribution of each

of these factors, additional GCM experiments with

separated forcings are planned. Here, we compare

instead our results with other model studies, which

differ in model type and prescribed forcing, and thus

allow to draw indirect conclusions on the relative

importance of the implied changes. It should however

be kept in mind that a detailed intercomparison of the

model studies would be required to distinguish between

the relative contributions of physical processes (e.g.,

solar irradiance fluctuations; changed atmospheric

chemistry due to composition changes; the ocean–atmo-

sphere feedback) and the impact of model and experi-

mental characteristics (e.g., representation of the

stratosphere; spectrally resolved versus total solar

irradiance changes; importance of interactive ozone

calculation; type of ocean model), which is beyond the

scope of this work.

There exists very good agreement in the regional

surface air temperature change patterns between this

study and the model results of Rind and Overpeck

(1993), Cubasch et al. (1997), Rind et al. (1999), and

Shindell et al. (2001): they all simulate a cooling of the

land masses and warming over the eastern North Pacific

and the western North Atlantic during MM. The

regional climate anomalies of our simulation, e.g., the

shift towards a low NAO index climatology, fits very

well to the results of Shindell et al. (2001), who

compared the MM period with pre-industrial condi-

tions. Their model response was not influenced by

industrial changes in atmospheric composition. The

decrease of solar insolation during MM therefore

appears to be a sufficient external forcing to produce

the described regional surface air temperature change

patterns. The near-surface temperature anomalies in

Shindell et al. (2001) are, however, about one-third of

those simulated in our study, indicating the additional

impact of the increase in industrial emissions between

MM and PD. In our MM simulation, the lower CO2

contributed to tropospheric cooling, and was the

dominant forcing for the stratospheric response.

In the FUB-CMAM, SSTs were adapted to the MM

solar forcing by prescribing SST anomalies from a

simulation of a coupled ocean atmosphere GCM

(Fischer-Bruns et al., 2002). The importance of the

ARTICLE IN PRESSU. Langematz et al. / Journal of Atmospheric and Solar-Terrestrial Physics 67 (2005) 55–69 67

ocean response to solar forcing was demonstrated by

Cubasch et al. (1997): in their GCM simulation, a delay

of 25 years in the near-surface temperature response to

an implied solar forcing was explained by a reduction of

the thermohaline circulation in the North Atlantic.

These results are supported by Fischer-Bruns et al.

(2002), who simulated a strong cooling of the surface air

temperature over the western North Atlantic, which they

explain by an increased sea-ice cover during the LMM

only to be captured in a transient simulation with a fully

coupled ocean model. Interestingly, both studies simu-

late temperature responses in this area of opposite sign.

So the inclusion of a full ocean model with thermohaline

circulation did not yield to a uniform temperature

response in both studies. In addition, our FUB-CMAM

simulation, using the Fischer-Bruns et al. (2002) SST

anomalies, simulates a warming, more similar to

Cubasch et al. (1997) than to Fischer-Bruns et al.

(2002), and also more similar to Shindell et al. (2001)

and Rind et al. (1999), who used simplified ocean

modules. So in summary, although the MM simulations

demonstrate the important role of the oceans, so far they

have not allowed to quantify the regional surface climate

response.

Another point of interest is the importance of the

stratosphere for the surface climate response. In our

simulation, we found an enhanced interaction between

the troposphere and the lower stratosphere as well as an

impact of the composition changes in the upper strato-

sphere. Shindell et al. (2001) pointed out the necessity to

include a highly resolved stratosphere in their model to

obtain the simulated shift in the variability modes for

decreased solar irradiance. Kodera (2002) recently

showed on the basis of observational data that the

NAO is influenced by the 11-year solar signal and

highlights the importance of the stratosphere. A more

detailed investigation of the downward impact of the

solar irradiance changes and their relative importance

for the surface climate anomalies in the MM simulation

will be the subject of future work.

One possible contribution to climate change during

MM has been neglected in this study: the impact of

volcanic eruptions. Crowley (2000) obtained a clear

surface cooling during the Little Ice Age caused by

volcanism in a linear energy balance model. In contrast,

Lean and Rind (1998) showed that NH summer

temperatures were highly correlated with the recon-

structed solar forcing ðr ¼ 0:86Þ between 1610 and 1800,

while they were only weakly correlated with the volcanic

dust veil index ðr ¼ �0:005Þ. In a recent GCM simula-

tion, both forcings, solar and volcanic, yield the best

agreement with historical reconstructions of global

mean temperature, however, regional changes were

dominated by solar forcing (Shindell et al., 2003). More

research is needed to clarify the role of volcanoes for

climate change during MM, and volcanic forcing will

therefore be included in our coming MM GCM

studies.

It should finally be noted that the focus of this study

was to examine the impact of known changes in external

forcing on the climate during MM. We could show that

the sum of the forcings led to a climate change in MM

which is very well in agreement with reconstructions

from proxy data. We could not show which of the

forcings contributes how much to the simulated changes.

This needs additional model simulations with separated

forcings. They are in progress and will supplement the

results shown here.

Acknowledgements

We wish to thank Judith Lean for solar irradiance

data, Joanna Haigh for ozone change data and Ulrich

Cubasch and Irene Fischer-Bruns for SST data. We are

especially grateful to Lon Hood for discussions during a

stay of UL at the Lunar and Planetary Laboratory,

University of Arizona. We further thank Gerard Bond

and an anonymous reviewer for their constructive

comments. This study was part of the project ‘‘Solar

Influence on Climate and the Environment (SOLICE,

EVK2-2002-00543)’’ supported by the European Com-

mission. The FUB-CMAM was run at the Konrad-

Zuse-Zentrum fur Informationstechnik, Berlin.

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