The role of stratosphere-troposphere coupling in the occurrence of

extreme winter cold spells over northern Europe

Lorenzo Tomassini,1 Edwin P. Gerber,2 Mark P. Baldwin,3 Felix Bunzel,1 and Marco Giorgetta1

Received 26 June 2012; accepted 26 August 2012; published 12 October 2012.

[1] Extreme cold spells over Northern Europe during winter are examined in order toaddress the question to what degree and in which ways stratospheric dynamics mayinfluence the state of the troposphere. The study is based on 500 years of a pre-industrial control simulation with a comprehensive global climate model which wellresolves the stratosphere, the MPI Earth System Model. Geopotential heightanomalies leading to cold air outbreaks leave imprints throughout the atmosphereincluding the middle and lower stratosphere. A significant connection betweentropospheric winter cold spells over Northern Europe and erosion of the stratosphericpolar vortex is detected up to 30 hPa. In about 40 percent of the cases, the extremecold spells are preceded by dynamical disturbances in the stratosphere. The strongwarmings associated with the deceleration of the stratospheric jet cause thetropopause height to decrease over high latitudes. The compression of thetropospheric column below favors the development of high pressure anomalies andblocking signatures over polar regions. This in turn leads to the advection of cold airtowards Northern Europe and the establishment of a negative annular mode patternin the troposphere. Anomalies in the residual mean meridional circulation during thestratospheric weak vortex events contribute to the warming of the lower stratosphere,but are not key in the mechanism through which the stratosphere impacts thetroposphere.

Citation: Tomassini, L., E. P. Gerber, M. P. Baldwin, F. Bunzel, and M. Giorgetta (2012), The role of stratosphere-

troposphere coupling in the occurrence of extreme winter cold spells over northern Europe, J. Adv. Model. Earth Syst., 4,

M00A03, doi:10.1029/2012MS000177.

1. Introduction

[2] Extended periods of cold temperatures in Europeduring wintertime are generally of dynamic origin. Highpressure anomalies over the Arctic or Siberia and lowpressure over the European continent induce the advec-tion of cold air from the north or the northeast towardslower latitudes. When this synoptic condition persistsover several days, a cold spell occurs.

[3] The described pattern of tropospheric pressureanomalies can sometimes be characterized by a negativephase of the North Atlantic Oscillation (NAO), as in thecold winter of 2010 [Jung et al., 2011]. In other instances,when the high pressure anomaly is centered more to theeast over Scandinavia or Siberia, the dynamic situationis better described as an atmospheric block and does not

well project on the NAO pattern. The anomalously coldwinter of 2006 represents an example for such a case[Croci-Maspoli and Davies, 2009].

[4] In both situations, the negative NAO phase andthe high latitude blocking, the general westerly flowregime over the North Atlantic region is disturbed.Such a disturbance can sometimes extend through thewhole troposphere and even into the stratosphere. Forinstance, during the cold winter of 2006, the northernhemispheric stratospheric vortex was particularly weak[Scaife and Knight, 2008; Jung et al., 2010].

[5] The stratospheric polar vortex is a feature of thewinter hemisphere climate. Although the stratosphericlarge-scale zonal flow is forced by differential diabaticheating, the winter stratosphere is not in radiativeequilibrium [Waugh and Polvani, 2010]. Verticallypropagating planetary waves from the troposphereperturb the stratospheric circulation resulting in zonalwinds that are weaker than predicted by radiativeequilibrium [Andrews et al., 1987]. Such waves occa-sionally lead to rapid deceleration of the zonal flowand accompanying sudden warmings of the stra-tosphere [Schoeberl, 1978].

[6] The traditional view on the coupling of the tro-posphere and the stratosphere is therefore centered on

1Max Planck Institute for Meteorology, Hamburg, Germany.2Center for Atmosphere Ocean Science, Courant Institute of

Mathematical Sciences, New York University, New York, NewYork, USA.

3College of Engineering, Mathematics and Physical Sciences,University of Exeter, Exeter, UK.

’2012. American Geophysical Union. All Rights Reserved.1942-2466/12/2012MS000177

JOURNAL OF ADVANCES IN MODELING EARTH SYSTEMS, VOL. 4, M00A03. doi:10.1029/2012MS000177, 2012

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the impact of upward propagating tropospheric waveson the stratospheric dynamics. Baldwin and Dunkerton[1999] however showed by composite analysis of vertic-ally resolved Northern Annular Mode (NAM) indicesthat stratospheric disturbances appear to propagatedownward and become manifest in changes to thetropospheric circulation. In Baldwin and Dunkerton[2001] it is argued that weak stratospheric vortexregimes during northern winter cause statistically sig-nificant changes in the probabilities of cold air out-breaks across Europe, Asia, and North America.

[7] However, since annular modes are calculated fromgeopotential height anomalies, they are partly an expres-sion of temperature anomalies in underlying atmo-spheric layers. A vertically coherent structure inannular mode variability is therefore implicit in theirdefinition. The central research question in this contextthus concerns the issue to what extent and throughwhich mechanisms a disturbance in the stratosphericpolar vortex may propagate downwards and affect thetropospheric circulation.

[8] The difficulty of this question consists in the factthat tropospheric disturbances may influence the flowregime in the troposphere as well as the dynamics of thestratosphere. Polvani and Waugh [2004] show that theinitial stratospheric NAM anomalies in Baldwin andDunkerton [2001] are themselves forced by wave break-ing events originating from the troposphere. In Gerberand Polvani [2009], idealized model experiments showthat topography can influence the position of the tro-pospheric jet as well as the variability in the stra-tospheric polar vortex. Thus, the correlation of thestrength of stratospheric vortex with the position ofthe tropospheric jet does not necessarily imply thatdisturbances in the stratospheric circulation cause thetropospheric jet to shift.

[9] A number of theories about the influence of thestratosphere on the troposphere have been proposed.Hartley et al. [1998] suggest that a redistribution ofstratospheric potential vorticity induces perturbationsin the upper troposphere [see also Black, 2002; Black andMcDaniel, 2004]. Ambaum and Hoskins [2002] stress therole of a change in tropopause height caused by stra-tospheric warmings [see also Cai and Ren, 2007]. Songand Robinson [2004] describe mediating feedbacks to thestratospheric forcing by transient eddies [see alsoKushner and Polvani, 2004; Chen and Held, 2007].Another line of reasoning concerns the modulation ofreflected upward propagating tropospheric planetarywaves by the strength of the wind shear in the lowerstratosphere [Perlwitz and Harnik, 2003].

[10] In the present paper we approach the question ofpossible impacts of stratospheric disturbances on thetropospheric circulation from the perspective of thetroposphere. Strong and persistent surface temperatureanomalies over Northern Europe are taken as a startingpoint of the investigation. This allows for studying caseswhere weak stratospheric polar vortex events and asso-ciated stratospheric warmings precede extreme negativetemperature anomalies in the troposphere. In suchinstances atmospheric fields show distinct perturbations

near the crucial region of the tropopause. By focusingon exceptionally strong signals, interactions of keydynamical features during stratosphere-tropospherecoupling can be identified more clearly in the otherwisevariable mid-latitude climate. The concentration onextreme events implies that long time series have to beconsidered. Therefore the study centers on a long pre-industrial control integration with a coupled climatemodel, the MPI Earth System Model, that well resolvesthe stratosphere.

[11] The article is structured as follows. In the secondsection the climate model and the simulation aredescribed. The definition of cold spells and their tro-pospheric characteristics are contained in Section 3.Section 4 examines the question whether the anomaliesin the stratosphere and the troposphere are related toeach other in a significant way. In Section 5 a classifica-tion of cold spells is given based on the extent of thevertical coupling that can be observed during or beforethe event. Several diagnostics that illuminate the phys-ical mechanisms acting during stratosphere-tropospherecoupling are presented for cases of downward propagat-ing disturbances in Section 6. Finally, a set of conclu-sions completes the paper.

2. Climate Model and Experiment

[12] The Max Planck Institute Earth System Model(MPI-ESM) is based on the atmospheric general circula-tion model ECHAM6 (B. Stevens et al., The atmosphericcomponent of the MPI Earth System Model: ECHAM6,manuscript in preparation, 2012), and an improvedversion of the Max Planck Institute Ocean Model MPI-OM (J. Jungclaus et al., Characteristics of the oceansimulations in the MPI Earth System Model, manuscriptin preparation, 2012). In the low resolution version(MPI-ESM-LR) used in this study, the atmosphericmodel utilizes a spectral transform dynamical core withtriangular truncation at wavenumber 63, associated witha horizontal Gaussian grid with resolution of about 1.9degrees. The vertical grid resolves the atmosphere in 47levels up to 0.01 hPa. The model includes the Hinesparameterization [Hines, 1997a, 1997b] for the upwardpropagation of unresolved non-orographic gravity wavesand their dissipation, providing tendencies in the hori-zontal winds. Its implementation follows Manzini et al.[2006]. The MPI-OM ocean model is a primitive equationmodel with a free surface and a mass flux boundarycondition for salinity. A simple bottom boundary layerscheme is included as well as the standard set of subgrid-scale parameterizations. The model possesses a typicalresolution of 1.5 degrees near the equator and 40 levels inthe vertical.

[13] The study focusses on 500 years of a MPI-ESM-LR pre-industrial control simulation with conditions ofthe year 1850, following the CMIP5 protocol [Tayloret al., 2012]. More details on the setup and performanceof the simulation can be found in M. Giorgetta et al.(Climate variability and climate change in the MPI-ESM CMIP5 simulations, manuscript in preparation,2012).

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3. Winter Cold Spells Over Northern Europe

[14] For the definition of winter cold spells overNorthern Europe we consider daily mean land temper-ature averaged over the geographic region of 0 East to40 East and 48 North to 65 North (see the red box inFigure 1, left). A cold spell is identified if the temper-ature drops below the 10 percent quantile of the cli-matological distribution for at least 15 consecutive days.In order to obtain a robust estimate of the 10 percentquantile, a 10 day moving window is used to calculatethe daily climatological temperature distribution foreach day in the year. Furthermore, we restrict ourattention to cold spells that start or end in December,January, or February.

[15] This definition of winter cold spells results in 41of such cold air outbreaks over Northern Europe duringthe 500 years of the MPI-ESM pre-industrial controlsimulation. A composite of the 2-meter temperatureanomaly for all winter cold spells is shown in Figure 1(left). In the composite the temperature anomaly israther confined over northeastern Europe. There is anextension towards Siberia but with an evidently flatten-ing amplitude in eastern direction. In contrast, a distinctwarm anomaly can be observed over Greenland, adja-cent regions of the Arctic Ocean, and the most easternpart of Siberia. This is in qualitative agreement withcharacteristics of recently observed European cold spells[e.g., Croci-Maspoli and Davies, 2009].

[16] The corresponding composite of geopotentialheight anomalies at 500 hPa (Figure 1, right) demon-strates the dynamic nature of the phenomenon. Lowpressure slightly south of the region of cold air temper-ature, and a high pressure anomaly over the highlatitudes north of Europe cause cold air from the north-east to advance towards the continent. The pattern isreminiscent of a negative phase of the North Atlantic

Oscillation (NAO), but the NAO has its usual center ofaction over the North Atlantic Ocean. In the case of thecold spells the dominant geopotential height anomaliesare located over land and north of Scandinavia.However, a trough of the low pressure area extends inwestward direction towards the East Coast of theUnited States, an indication that in some cases a nega-tive phase of the NAO is indeed responsible for the coldtemperatures over Northern Europe.

[17] A comparison of the simulated distribution ofmonthly mean winter land temperature anomalies overNorthern Europe with observations [Casty et al., 2007]shows that the pre-industrial control run well reproduceswintertime variability in this region (Figure 2). The lowertail of the distribution is particularly well captured by themodel. Also stratospheric variability and the coupling ofthe stratosphere to the troposphere are satisfactorilyrepresented by the model (A. J. Charlton-Perez et al.,Mean climate and variability of the stratosphere in theCMIP5 models, manuscript in preparation, 2012), as wellas general Northern Hemispheric climate features(Stevens et al., manuscript in preparation, 2012).

[18] The most prominent feature of the geopotentialheight anomaly composite in Figure 1 consists in thehigh latitude block. In some cold spell events it issituated right north of the continent, in some instancesit is displaced slightly to the east or to the west. We use atwo-dimensional version of the Tibaldi-Molteni [Tibaldiand Molteni, 1990] blocking index as defined in Scherreret al. [2006] in order to quantify the occurrence of highpressure blocks during the cold spells. The index assignsa 1 to days which show a blocking signature, and 0 todays without.

[19] Figure 3 (left) shows the winter (NDJFM) block-ing climatology of the pre-industrial control run, theright panel the composite of the blocking occurrenceover the period of all cold spells subtracted from the

Figure 1. (left) Composite of the 2-meter temperature anomaly for all winter cold spells. (right) Correspondingcomposite of geopotential height anomalies at 500 hPa.

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climatology. In the composite the center of the block islocated over the northern tip of Scandinavia stretchingtowards the east over the Barents Sea. This is in goodagreement with studies of winter time blocking and theireffect on the tropospheric climate based on reanalyses[e.g., Trigo et al., 2004].

4. Significance of Stratosphere-TroposphereInteraction

[20] In the present paper we investigate the role ofstratosphere-troposphere interactions in the devel-opment of Northern European winter cold spells.

Since the seminal paper by Baldwin and Dunkerton[2001] the Northern Annual Mode (NAM) time seriesserves as most prominent stratosphere-tropospherecoupling index [see also Baldwin and Dunkerton, 1999].In the following we use a NAM index based on dailyzonally-averaged geopotential height as recommendedin Baldwin and Thompson [2009] and employed also byGerber et al. [2010]. The zonal-mean NAM is less proneto contamination by the Pacific/North American tele-connection (PNA) pattern, reflects well the daily evolu-tion of stratosphere-troposphere coupling events, andcan be computed consistently over the 500 years of thepre-industrial control simulation considered here.

Figure 2. Distribution of monthly winter temperature anomalies over the land area of 48 North to 65 North and 0East to 40 East based on observations for the period 1765 to 1964 [Casty et al., 2007] and 500 years of the MPI-ESM-LR pre-industrial control simulation.

Figure 3. (left) Winter (NDJFM) blocking climatology of the MPI-ESM pre-industrial control run. (right)Composite of the blocking occurrence anomalies over the period of all cold spells. Blocking occurrence is measuredin percentage of days with blocking signature.

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[21] In the stratosphere the zonal mean NAM indexstrongly correlates with zonal-mean wind anomaliesbetween 40 North and 90 North [Baldwin andThompson, 2009]. Accordingly, for the present section,we define weak vortex events formally by requiring thecentered NAM index to be below -1 times its standarddeviation for at least 15 consecutive days. The persist-ence criterion is motivated by the interest in the relationof weak vortex events to cold spells. The NAM indextime series is slightly smoothed using a 6-day Lanczosfilter before the calculation of the weak vortex events inorder not to let short nonfulfillments of the criterionhave a major impact. Although in the troposphere thethus defined weak vortex events would more appropri-ately be called negative NAM events, for simplicity wewill not make this terminological distinction.

[22] Here we would like to assess if the connectionbetween winter cold spells over Northern Europe andpreceding weak vortex events is significant. This requiresthe definition of a significance measure and, as a pre-requisite, a criterion that selects weak vortex events thatare related to cold spells, at least in time.

[23] Regarding the last point we define a weak vortexevent to coincide with a cold spell if the weak vortexevent starts before the end of the cold spell and ends nosooner than 40 days before the start of the cold spell.For the stratosphere higher than 100 hPa, ‘‘40 days’’ isreplaced by ‘‘60 days’’ in the condition. This is moti-vated by Baldwin and Dunkerton [2001] where it is shownthat the downward propagation of disturbances fromthe upper stratosphere to the lower troposphere exhibitsa timescale of up to about 2 months.

[24] In the following we assess the significance of thenumber of weak vortex events that coincide, in theabove defined sense, with winter cold spells. To thisend, the number of coincidences have to be related to thefrequency of occurrence of weak vortex events asimplied by their definition.

[25] For each month in the year the frequency ofoccurrence of weak vortex events is first diagnosed fromthe pre-industrial control run. Then artificial start datesof weak vortex events are randomly sampled from alldates of each calender month in such a way that for eachmonth in the year the number of such artificially gener-ated weak vortex event start dates is the same as in thepre-industrial control run. Analogously, correspondingartificial end dates are determined in such a way that themean and the standard deviation of the length of theartificially generated weak vortex events is the same asthe mean and the standard deviation of the length of theactual weak vortex events in the pre-industrial controlsimulation.

[26] Thus, the synthetically generated weak vortexevents are equally frequent, but decorrelated in timefrom the simulated cold spells. This process is repeated109000 times. The distribution of coincidences of theartificial weak vortex events (Figure 4) can be comparedto the actual number of coincidences in the pre-indus-trial control run (red lines in Figure 4). If the actualnumber of coincidences is at the upper end of thedistribution of random coincidences, it is likely that

there is a physical mechanism linking cold spells to weakvortex events in the pre-industrial control run. The greyshaded areas in Figure 4 indicate 2 standard deviationsof the distributions.

[27] Between 100 hPa and the surface, the considera-tion of a 40 days period and a 60 days period before thecold spells give qualitatively similar results, while higherup in the stratosphere the 40 days condition clearly leadsto a less significant relation between weak vortex eventsand cold spells. Less significance in the upper stra-tosphere is not caused by lower numbers of actualcoincidences in the control simulation, but by longercorrelation times in the NAM index. The latter factimplies that the persistence criterion in the definition ofweak vortex events is more readily met, which leads toan increase of the frequency of weak vortex events.

[28] The significance of the relation of weak vortexevents and cold spells that is evident in the troposphereand the lower stratosphere motivates to further invest-igate the physical mechanisms that vertically couple thestratosphere and the troposphere and relate surface coldspells to preceding stratospheric weak vortex events.

5. Dynamical Classification of Cold Spells

[29] As can be seen from Figure 5, the geopotentialheight anomalies over Northern Europe during theperiod of all cold spells persist up to the stratosphereand change their characteristics only at the 10 hPa level.This is not per se particularly surprising as the geopo-tential height field tends to integrate temperature anom-alies in the vertical.

[30] The main point of interest in the present studytherefore is the question whether there are cases of coldspells in which the origin of the dynamical disturbancecan be traced back to the stratosphere. To examine thisissue in more detail we categorize the 41 cold spells inthe pre-industrial control run.

[31] We call a cold spell stratospherically induced, ordownward propagating, if we detect a coinciding weakvortex event at the 10 hPa as well as at the 100 hPa level.Here the definition of coinciding weak vortex events ofSection 4 is adopted with two slight modifications. At100 hPa, we require the NAM index to stay below minus0.8 times its standard deviation (instead of minus 1 timesits standard deviation) to account for the higher vari-ability at lower levels of the stratosphere. Moreover, theweak vortex events need to start strictly before the startof the cold spells.

[32] Cold spells which meet this criterion at 100 hPabut not at 10 hPa are referred to as lower stratosphericcases. Finally we also need to formulate a formaldefinition of cold spells that are dynamically restrictedto the troposphere. As can be seen from Figure 5, thedynamical characteristics of the northern Europeanwinter cold spells exhibit distinct regional features whichdo not always project equally well on the NAM patternin the troposphere. Therefore, in order to detect atmo-spheric flow anomalies in the troposphere related toEuropean cold spells, we introduce a more regional andflexible geopotential height based index than the NAMindex.

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[33] This regional geopotential height index is definedby subtracting the maximum of the geopotential heightanomaly over the area 30W to 90E and 65N to 80N

from the minimum of the geopotential height anomalyover the area 0E to 40E and 45N to 55N. Applying thisdefinition to the daily geopotential height anomaly field,

Figure 4. Distribution of coincidences of artificially sampled weak vortex events with the winter cold spells atdifferent pressure levels. The actual number of coincidences in the pre-industrial control run are indicated by redlines. Grey shaded areas indicate 2 standard deviations of the distributions.

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and normalizing by the mean and standard deviation,results in a daily regional geopotential height index.

[34] As can be seen from the geopotential heightcomposites in Figure 5, negative anomalies are prevalentover Europe between 45N to 55N, and positive anom-alies dominate north of the cold spell area in a band thatcovers 30W to 90E from the lower troposphere up to30 hPa. Using maxima and minima in the definitionmakes the index less dependent on a specific pre-selected

pattern as in the case of the NAM. This flexibility isuseful since the individual events have somewhat differ-ent synoptic features. As the regional north-south dipoleof geopotential height anomalies, common to all cases,determines the specific atmospheric flow conditionwhich leads to the cold spells, the index is suitable todescribe the dynamical nature of the cold air outbreaks.

[35] Cold spells that do not comply with the criteria ofstratospherically induced or lower stratospheric events,

Figure 5. Geopotential height anomalies over the Northern Hemisphere during the period of all cold spells atdifferent pressure levels.

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but for which the regional geopotential height indexstays below minus 0.8 times its standard deviation for atleast 15 consecutive days at the 500 hPa level within aperiod of 40 days before the start of the cold spell andthe end of the cold spell, are termed tropospheric cases.

[36] Of the 41 cold spells there are 17 stratosphericallyinduced events, 8 lower stratospheric events, and 15tropospheric cases. One cold spell does not fall in eitherof the categories. It is characterized by a strong andextended negative pressure anomaly over the Europeancontinent, Siberia, and large parts of the Arctic, andprojects on a positive NAM index. Dynamically it ismainly confined to the troposphere, and we will notinclude this event in the following composite analysis.

[37] In order to discuss the dynamical signatures ofthe three classes we introduce two more indices similarin spirit to the regional geopotential height index. Aregional wind index is defined as the mean of the zonalwind anomaly over Northern Europe where the coldspells occur (0E to 40E and 48N to 65N). In thetroposphere it is indicative of the easterly flow whichleads to the cold air outbreaks, while in the stratosphereit characterizes the strength of the polar vortex.Analogously, a regional temperature index is explainedas the mean of the temperature anomalies over the area0E to 40E and 55N to 65N. Both indices are normalizedby their respective means and standard deviations. Thetemperature index is confined to 55N (instead of 48N)since at higher levels of the atmosphere, temperatureanomalies that are associated with dynamical distur-bances in the stratosphere tend to weaken towards lowerlatitudes [Schoeberl, 1978].

[38] Figure 6 shows the time development of thedifferent indices for the three cold spell classes. Day 0in the stratospherically induced case is defined as thestart of the weak vortex event at the 10 hPa level. Forthe lower stratospheric case, day 0 in the composite isdetermined by the start of the weak vortex event at the100 hPa level. In the tropospheric case day 0 corre-sponds to the start of the dynamical disturbance asdefined by the regional geopotential height index at500hPa.

[39] The panels in Figure 6 include stippled areaswhich indicate a measure of significance with respectto natural variability. For each index and level of theatmosphere, N arbitrary winter days are sampled, whereN is the number of events that enter the respectivecomposite. The mean is calculated over these N days.This procedure is then repeated 109000 times. Thestippled regions indicate the areas where the compositeexceeds 2 standard deviations of the distributions ofthese mean values.

[40] One can see that the tropospheric events areindeed mainly restricted to the troposphere. Duringthe cold spell one can identify a signal in the geopoten-tial height index (and, in accordance, in the zonal windand the temperature index) also in the stratosphere, butthis is rather a consequence of the temperature anomalyat lower levels than the cause of the cold spell.

[41] Also for the lower stratospheric cases there areindications that the dynamical disturbances originate in

the troposphere. In all indices there are, to some degree,signals in the troposphere before day 0 of the composite,and effects in the stratosphere occur instantaneously orlater in time.

[42] The situation is different, although not comple-tely unambiguous, in the case of the stratosphericallyinduced events. In the NAM index the signal in thestratosphere shows up prior to the disturbance in thetroposphere. It is evident rather instantaneouslythroughout the stratosphere, but the persistence of theanomalies is larger in the lower stratosphere. However,in the regional geopotential height index and theregional wind index, anomalies are present in the tro-posphere before the main development of the events inthe stratosphere. The tropospheric disturbances couldhave been, at least partially, the cause of increased wavebreaking and deceleration of the vortex in the stra-tosphere [Cohen et al., 2007]. Black and McDaniel[2004] give evidence that the possibility for stra-tosphere-troposphere interactions depends on the pre-existing state of the troposphere. A particularly cleardownward propagating signal seems to be present in theregional mid-latitudinal temperature index. In the nextsection we focus on the stratospherically induced coldspell events and further investigate the physical mechan-isms of the stratosphere-troposphere coupling.

6. Mechanisms of Stratosphere-TroposphereCoupling

[43] When the composite of temperature anomaliesover Northern Europe (first row of Figure 7, left, sameas in Figure 6) is compared to the temperature anom-alies calculated over the polar area of 0E to 40E and70N to 90N (first row of Figure 7, right), it is apparentthat over the polar region the positive temperatureanomalies become effective immediately after the startof the vortex disturbance across the whole stratosphere.As a consequence, the tropopause is continuously low-ered (fourth row of Figure 7). Here again an areaaverage is taken over the region 0E to 40E and 70N to90N for the right panel, and 0E to 40E and 55N to 65Nfor the Figure 7 (left).

[44] With some time delay negative temperatureanomalies, indicative of changes in the troposphericcirculation, develop over Northern Europe, and nega-tive potential vorticity anomalies accumulate in thelower stratosphere and penetrate through the tro-popause, mostly over the polar region (third row ofFigure 7).

[45] Here we use the thermal definition of the tro-popause according to which the tropopause is identifiedas the lowest altitude where the temperature lapse rate issmaller than 2 Kelvin per km, provided that the averagelapse rate from this level to any point within 2 km abovealso features a lapse rate smaller than 2 Kelvin per km(see Gettelman et al. [2011] for a discussion of differenttropopause definitions). The change in the tropopauseheight during weak vortex events is an immediate con-sequence of the strong warming of the lower stra-tosphere, and is further amplified by dynamicalfeedbacks which cause a cooling in the troposphere.

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[46] An impact of the strength of the polar vortex onthe tropospheric circulation is established in Ambaumand Hoskins [2002] by exploring the effect of potentialvorticity anomalies in the stratosphere on the tro-popause height, and the consequence of changed tro-popause height for surface pressure over the North Pole.The theory is developed in the context of a quasigeos-trophic model from which a potential vorticity equationis derived. Here again, negative potential vorticityanomalies in the stratosphere result in a lowered tro-popause, a compressed tropospheric column below, andconsequently a reduced relative vorticity over the polar

cap associated with a high pressure signal. Similar ideaswere put forward in Hartley et al. [1998] and Black[2002], and a general discussion of what factors controlthe tropopause height is contained in Schneider [2007].Black and McDaniel [2004] point out that the potentialvorticity anomalies need to descend to sufficiently lowaltitudes within the stratosphere in order to have aneffect on the troposphere.

[47] Figure 8 presents the time development of poten-tial vorticity anomalies before and during the downwardpropagating cold spells, significant regions are stippled.Figure 8 (left) contains a mean over days 30 to 15 before

Figure 6. Time development of different indices for (left) stratospherically induced, (middle) lower stratosphericand (right) tropospheric cold spells: (first row) NAM index, (second row) regional geopotential height index, (thirdrow) regional wind index, and (fourth row) regional temperature index. Stippling indicates areas where theanomalies exceed 2 standard deviations of natural variability.

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Figure 7. Time development of different diagnostics for stratospherically induced cold spells. (first row) Mid-latitudinal temperature anomalies and polar temperature anomalies. (second row) Mid-latitudinal E-P fluxdivergence anomalies and mid-latitudinal residual circulation anomalies. (third row) Mid-latitudinal potentialvorticity anomalies and polar potential vorticity anomalies. (fourth row) Mid-latitudinal tropopause heightanomalies and polar tropopause height anomalies.

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Figure 8. Time development of potential vorticity anomalies before and during the stratospherically induced coldspells. (left) Mean over days 30 to 15 before. (middle) Mean over day 15 to day 1 before. (right) Mean over day 1 today 15 of the cold spells. Areas that exceed 2 standard deviations of natural variability are stippled.

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the start of the cold spells, Figure 8 (middle) a mean overday 15 to day 1 before, and Figure 8 (right) a mean overday 1 to day 15 of the cold spells. One can observe thatnegative potential vorticity anomalies develop at 50 hPa(first row of Figure 8) over polar regions before the coldspells. They intensify during the 15 days before the coldspells, and become evident in the lower stratosphere aswell. In the troposphere, at the 500 hPa geopotentialheight level (fourth row of Figure 8), the anomalies aremost pronounced during the cold spells.

[48] Polvani and Kushner [2002] report, based onsimulations with a simplified general circulation model,a poleward shift of the tropospheric jet when the stra-tosphere is cooled and thus the stratospheric vortexstrengthened. Similarly, Williams [2006] describes howincreased temperatures in the stratosphere and an asso-ciated lower tropopause height shift the tropospheric jetequatorwards in idealized model simulations.

[49] This mechanism implies a certain time delay andpotentially an elongation of time scales. Gerber andVallis [2007] show that in idealized model experimentsthe interaction between synoptic eddies retard themotion of the jet, slowing its meridional variation andthereby extending the persistence of annular modeanomalies in the troposphere. The negative zonal-meanzonal wind anomalies are evident in the stratosphere in

the days 45 to 30 before the start of the cold spells(Figure 9, top right). During day 30 to day 1 before thecold spells (Figures 9, bottom left, and 9, bottom mid-dle), the poleward flank of the tropospheric jet over theNorth Atlantic gradually weakens, while the equator-ward side strengthens. The negative zonal wind anom-alies of the stratosphere extend downward into thetroposphere over the region of the occurrence of the coldspells. The strengthening of the zonal wind at the south-ern side of the area of surface temperature anomalies isconsistent with the geopotential height anomaly pattern(Figure 5).

[50] Thus, a dynamical feedback can be identified.The lowering of the tropopause caused by the stra-tospheric warming over the polar cap produces a tro-pospheric high pressure anomaly over the polar regionwhich leads to the advection of cold air towardsNorthern Europe. The resulting cold air outbreakinduces negative geopotential height anomalies overthe continent which amplifies the negative NAM sig-nature. At the southern flank of the cold spell area themeridional pressure gradient and consequently the zonalflow are enhanced. In accordance with Charlton et al.[2005], the impact of the stratosphere on the troposphereis therefore mediated through synoptic-scale systemsand can not be understood merely as a large-scale

Figure 9. (top left) Winter (NDJFM) zonal-mean zonal wind climatology over the North-Atlantic sector (60 Westto 20 East). (top right) Zonal-mean zonal wind anomalies for day 45 to day 30 before the cold spells. Zonal-meanzonal wind anomalies for (bottom left) day 30 to day 15 before, (bottom middle) day 15 to day 1 before the coldspells, and (bottom right) day 1 to day 15 during the cold spells.

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adjustment of the troposphere to stratospheric potentialvorticity anomalies.

[51] Near the extratropical tropopause and lowerstratosphere, the stratospheric residual circulation pro-vides adiabatic warming which contributes to the low-ering of the extratropical tropopause [Son et al., 2007;Birner, 2010]. In the second row of Figure 7 a compositeof the Eliassen-Palm flux divergence (Figure 7, left) andthe residual circulation mass stream function (Figure 7,right) for the downward propagating cold spells areshown. Quantities are averaged over the latitudes 65Nto 75N, the qualitative picture does however not dependon the exact region considered. Increased Eliassen-Palmflux divergence during the early stage of the stra-tospheric disturbance confirms the crucial role of wavebreaking in the development of stratospheric weakvortex events and associated warmings [Schoeberl,1978]. Significant anomalies in the strength of theresidual circulation occur mainly at the beginning ofthe stratospheric vortex disturbance. This suggests thatthe residual mean meridional circulation does notdirectly play a decisive role in the mechanism throughwhich the stratosphere impacts the troposphere. Rather,in accordance with Cai and Ren [2007], the apparentdownward propagation of stratospheric disturbancesinto the troposphere is a consequence of the dynamicresponse to heating and cooling anomalies.

7. Conclusions

[52] Extreme winter cold spells over Northern Europewith a return period of about 12 years are investigated ina long pre-industrial control simulation using the MPIEarth System Model. A significant relation betweensuch cold air outbreaks and preceding circulation anom-alies in the stratosphere up to at least 50 hPa can bedetected. 17 out of 41 cold spells occur in associationwith a downward propagating dynamical disturbancewhich originates in the stratosphere. However, also inthese cases preexisting geopotential height anomaliesreminiscent of a negative annular mode pattern arepresent in the troposphere.

[53] A composite analysis of the downward propagat-ing cases reveals the important role of wave breaking inthe development of stratospheric vortex anomalies.Vertically propagating planetary waves disturb the stra-tospheric circulation and transport heat and momentumfrom mid-latitudes into the polar region, causing strongstratospheric warmings. Immediate temperature increasescan be observed over the polar cap, while over the mid-latitudes the temperature signal exhibits a downwardpropagating structure representing the dynamical evolu-tion of the stratospheric vortex disturbance. In the lowerstratosphere and near the tropopause, the temperatureanomalies persist due to the relatively low efficiency ofradiative cooling at these height levels [Kiehl and Solomon,1986]. Anomalies in the residual mean meridional circula-tion partly contribute to the adiabatic warming in thelower stratosphere, but the anomaly in the residualcirculation mass stream function is mainly restricted tothe period of the strongest dynamical disturbance in thestratosphere.

[54] As a consequence of the warming in the lowerstratosphere, the tropopause height is lowered whichleads to a compression of the tropospheric columnbelow and a high pressure signature over the polarregion in the troposphere. Zonal wind anomalies arethus transferred from the stratosphere to the tro-posphere, and the advection of cold air from the north-east towards Northern Europe further reinforces thenegative NAM pattern in the troposphere.

[55] The impact of the stratosphere on the tro-posphere can thus be understood as dynamical responseto heating anomalies rather than the effect of stra-tospheric momentum forcing on the meridional circula-tion as expressed in the so-called ‘‘principle ofdownward control’’ [Haynes et al., 1991], originallyformulated in a zonally symmetric context. The dynam-ical feedback amplifies the effect of the lower stra-tospheric perturbation in the troposphere.

[56] Acknowledgments. Useful comments by Elisa Manzini greatlyhelped to guide the analysis. Support by James Anstey, who providedthe code for the blocking index computation, as well as the Max PlanckSociety for the Advancement of Science, is gratefully acknowledged. Wethank Hauke Schmidt for a careful reading of the paper, and theGerman Climate Computing Center (DKRZ) for the provision of thesimulation. The work was partially funded by the EuropeanCommission’s 7th Framework Programme, under GA 226520 for theCOMBINE project.

ReferencesAmbaum, M. H. P., and B. J. Hoskins (2002), The NAO troposphere-

stratosphere connection, J. Clim., 15, 1969–1978, doi:10.1175/1520-0442(2002)015,1969:TNTSC.2.0.CO;2.

Andrews, D. G., J. R. Holton, and C. B. Leovy (1987), MiddleAtmosphere Dynamics, 489 pp., Academic, San Diego, Calif.

Baldwin, M. P., and T. J. Dunkerton (1999), Propagation of the ArcticOscillation from the stratosphere to the troposphere, J. Geophys.Res., 04, 30,937–30,946.

Baldwin, M. P., and T. J. Dunkerton (2001), Stratospheric harbingersof anomalous weather regimes, Science, 294, 581–584, doi:10.1126/science.1063315.

Baldwin, M. P., and D. W. J. Thompson (2009), A critical comparisonof stratosphere-troposphere coupling indices, Q. J. R. Meteorol.Soc., 135, 1661–1672, doi:10.1002/qj.479.

Birner, T. (2010), Residual circulation and tropopause structure,J. Atmos. Sci., 67, 2582–2600, doi:10.1175/2010JAS3287.1.

Black, R. X. (2002), Stratospheric forcing of surface climate in theArctic Oscillation, J. Clim., 15, 268–277, doi:10.1175/1520-0442(2002)015,0268:SFOSCI.2.0.CO;2.

Black, R. X., and B. A. McDaniel (2004), Diagnostic case studies ofthe Northern Annular Mode, J. Clim., 17, 3990–4004, doi:10.1175/1520-0442(2004)017,3990:DCSOTN.2.0.CO;2.

Cai, M., and R.-C. Ren (2007), Meridional and downward propaga-tion of atmospheric circulation anomalies. Part I: NorthernHemisphere cold season variability, J. Atmos. Sci., 64, 1880–1901,doi:10.1175/JAS3922.1.

Casty, C., C. C. Raible, T. F. Stocker, J. Luterbacher, and H. Wanner(2007), A European pattern climatology 1766–2000, Clim. Dyn., 29,791–805, doi: 10.1007/s00382-007-0257-6.

Charlton, A. J., A. O’Neill, P. Berrisford, and W. A. Lahoz (2005),Can the dynamical impact of the stratosphere on the troposphere bedescribed by large-scale adjustment of the stratospheric PV distri-bution, Q. J. R. Meteorol. Soc., 131, 525–543, doi:10.1256/qj.03.222.

Chen, G., and I. M. Held (2007), Phase speed spectra and the recentpoleward shift of Southern Hemisphere surface westerlies, Geophys.Res. Lett., 34, L21805, doi:10.1029/2007GL031200.

Cohen, J., M. Barlow, P. J. Kushner, and K. Saito (2007),Stratosphere-troposphere coupling and links with Eurasia landsurface variability, J. Clim., 20, 5335–5343, doi:10.1175/2007JCLI1725.1.

TOMASSINI ET AL.: STRATOSPHERE-TROPOSPHERE COUPLINGM00A03 M00A03

13 of 14

Croci-Maspoli, M., and H. C. Davies (2009), Key daynamical featuresof the 2005/2006 European winter, Mon. Weather Rev., 137, 664–678, doi:10.1175/2008MWR2533.1.

Gerber, E. P., and L. Polvani (2009), Stratosphere-troposphere coup-ling in a relatively simple AGCM: The importance of stratosphericvariability, J. Clim., 22, 1920–1933, doi:10.1175/2008JCLI2548.1.

Gerber, E. P., and G. K. Vallis (2007), Eddy-zonal flow interactionsand the persistence of the zonal index, J. Atmos. Sci., 64, 3296–3311,doi:10.1175/JAS4006.1.

Gerber, E. P., et al. (2010), Stratosphere-troposphere coupling andannular mode variability in chemistry-climate models, J. Geophys.Res., 115, D00M06, doi:10.1029/2009JD013770.

Gettelman, A., P. Hoor, L. L. Pan, W. J. Randel, M. I. Hegglin, andT. Birner (2011), The extratropical upper troposphere and lowerstratosphere, Rev. Geophys., 49, RG3003, doi:10.1029/2011RG000355.

Hartley, D. E., J. T. Villarin, R. X. Black, and C. A. Davis (1998),A new perspective on the dynamical link between the stratosphereand troposphere, Nature, 391, 471–474, doi:10.1038/35112.

Haynes, P. H., C. J. Marks, M. E. McIntyre, T. G. Shepherd, andK. P. Shine (1991), On the ‘‘downward control’’ of extratropicaldiabatic circulations by eddy-induced mean zonal forces, J. Atmos. Sci.,48, 651–678, doi:10.1175/1520-0469(1991)048,0651:OTCOED.2.0.CO;2.

Hines, C. O. (1997a), Doppler spread parameterization of gravity wavemomentum deposition in the middle atmosphere. Part I: Basicformulation, J. Atmos. Sol. Terr. Phys., 59, 371–386, doi:10.1016/S1364-6826(96)00079-X.

Hines, C. O. (1997b), Doppler spread parameterization of gravity wavemomentum deposition in the middle atmosphere. Part II: Broad andquasimonochromatic spectra and implementation, J. Atmos. Sol.Terr. Phys., 59, 387–400, doi:10.1016/S1364-6826(96)00080-6.

Jung, T., T. N. Palmer, M. J. Rodwell, and S. Serrar (2010),Understanding the anomalously cold European winter of 2005/06using relaxation experiments, Mon. Weather Rev., 138, 3157–3174,doi:10.1175/2010MWR3258.1.

Jung, T., F. Vitart, L. Ferranti, and J.-J. Morcrette (2011), Origin andpredictability of the extreme negative NAO winter of 2009/10,Geophys. Res. Lett., 38, L07701, doi:10.1029/2011GL046786.

Kiehl, J. T., and S. Solomon (1986), On radiative balance of thestratosphere, J. Atmos. Sci., 43, 1525–1534, doi:10.1175/1520-0469(1986)043,1525:OTRBOT.2.0.CO;2.

Kushner, P. J., and L. Polvani (2004), Stratosphere-tropospherecoupling in a relatively simple AGCM: The role of eddies, J.Clim., 17, 629–639, doi:10.1175/1520-0442(2004)017,0629:SCIARS.2.0.CO;2.

Manzini, E., M. A. Giorgetta, M. Esch, L. Kornblueh, and E. Roeckner(2006), The influence of sea surface temperatures on the northernwinter stratosphere: Ensemble simulations with the MAECHAM5model, J. Clim., 19, 3863–3881, doi:10.1175/JCLI3826.1.

Perlwitz, J., and N. Harnik (2003), Observational evidence of astratospheric influence on the troposphere by planetary wave

reflection, J. Clim., 16, 3011–3026, doi:10.1175/1520-0442(2003)016,3011:OEOASI.2.0.CO;2.

Polvani, L. M., and P. J. Kushner (2002), Tropospheric response tostratospheric perturbations in a relatively simple general circulationmodel, Geophys. Res. Lett., 29(7), 1114, doi:10.1029/2001GL014284.

Polvani, L. M., and D. W. Waugh (2004), Upward wave activity fluxas a precursor to extreme stratospheric events and subsequentanomalous surface weather regimes, J. Clim., 17, 3548–3554,doi:10.1175/1520-0442(2004)017,3548:UWAFAA.2.0.CO;2.

Scaife, A. A., and J. R. Knight (2008), Ensemble simulations of thecold European winter of 2005–2006, Q. J. R. Meteorol. Soc., 134,1647–1659, doi:10.1002/qj.312.

Scherrer, S. C., M. Croci-Maspoli, C. Schwierz, and C. Appenzeller(2006), Two-dimensional indices of atmospheric blocking and theirstatistical relationship with winter climate patterns in the Euro-Atlantic region, Int. J. Climatol., 26, 233–249, doi:10.1002/joc.1250.

Schneider, T. (2007), The thermal stratification of the extratropicaltroposphere, in The Global Circulation of the Atmosphere, edited byT. Schneider and A. H. Sobel, pp. 47–77, Princeton Univ. Press,Princeton, N. J.

Schoeberl, M. R. (1978), Stratospheric warmings: Observations andtheory, Rev. Geophys., 16, 521–538, doi:10.1029/RG016i004p00521.

Son, S.-W., S. Lee, and S. B. Feldstein (2007), Intraseasonal variabilityof the zonal-mean extratropical tropopause height, J. Atmos. Sci.,64, 608–620, doi:10.1175/JAS3855.1.

Song, Y., and W. A. Robinson (2004), Dynamical mechanisms forstratospheric influences on the troposphere, J. Atmos. Sci., 61, 1711–1725, doi:10.1175/1520-0469(2004)061,1711:DMFSIO.2.0.CO;2.

Taylor, K. E., R. J. Stouffer, and G. A. Meehl (2012), An overview ofCMIP5 and the experiment design, Bull. Am. Meteorol. Soc., 93,485–498, doi:10.1175/BAMS-D-11-00094.1.

Tibaldi, S., and F. Molteni (1990), On the operational predictability ofblocking, Tellus, Ser. A, 42, 343–365, doi:10.1034/j.1600-0870.1990.t01-2-00003.x.

Trigo, R. M., I. F. Trigo, C. C. DaCamara, and T. J. Osborn (2004),Climate impact of the European winter blocking epsiodes from theNCEP/NCAR reanalyses, Clim. Dyn., 23, 17–28, doi:10.1007/s00382-004-0410-4.

Waugh, D. W., and L. M. Polvani (2010), Stratospheric polar vortices,in The Stratosphere: Dynamics, Transport, and Chemistry, Geophys.Monogr. Ser., vol. 190, edited by L. M. Polvani, A. H. Sobel, and D.W. Waugh, pp. 43–57, AGU, Washington, D. C., doi:10.1029/2009GM000887.

Williams, G. P. (2006), Circulation sensitivity to tropopause height,J. Atmos. Sci., 63, 1954–1961, doi:10.1175/JAS3762.1.

Corresponding author: L. Tomassini, Max Planck Institute forMeteorology, Bundesstrasse 53, D-20146 Hamburg, Germany.(lorenzo.tomassini@zmaw.de)

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