the climate during the maunder minimum: a simulation with the freie universität berlin climate...
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(U. Langematz
Journal of Atmospheric and Solar-Terrestrial Physics 67 (2005) 55–69
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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 maximaof 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 budgetrevealed 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 NHand 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 duringNH 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 anenhanced 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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