Un 2ième réchauffement stratosphérique (RSS²) a été observé  avec le maximum atteint le 6 janvier 2019.


Dernier update:   16 janvier 2019 11h00

Prévision des températures à 2m en Belgique (Modèle GFS/USA) (run 15/1/19 06z)

Prévisions de l'anomalie thermique à 2m (écart à la température moyenne) pour la période du 16/1/19 au 27/2/19 Run du 13-16/1/19 12z (MODELE CFSv2/USA)

Prévisions du centre européen (EPS-ECMWF): du 23 au 27 janvier  2019 (run du 13/1/19 00z)

Graphiques reprenant la modification du profil du vent sur le 65 parallèle Nord à 30 km d'altitude (pression atmosphérique = 10 hPa).

Les valeurs positives mentionnent la vitesse moyenne des vents ayant une composante "ouest" et les valeurs négatives la vitesse moyenne des vents ayant une composante "est".


(La prévision est en rouge)



Publié précédemment:

Un 2ième réchauffement stratosphérique (RSS²) a été observé  avec le maximum atteint le 6 janvier 2019.

Dates à retenir concernant ce deuxième réchauffement stratosphérique soudain (RSS²)

en tenant compte du temps de réponse moyen (+/- 14 jours) pour l'impact potentiel du réchauffement stratosphérique et du splitting entrainant des bouleversements des courants dans la stratosphère et ensuite dans la troposphère:

1. Le réchauffement stratosphérique maximum a eu lieu le 6/1/19 + +/-14 jours = +/-20 janvier 2019

2. Le split du vortex polaire dans haute stratosphère (10 hPa) aurait à nouveau lieu le 11/1/19

3. Maximum du vent zonal moyen de composante EST sur le 60°N dans la haute stratosphère (10hPa) :

Le 10 janvier 2019 , le vent moyen de composante EST sur le 60N a atteint son maximum de -10.7m/s .

Il est prévu que le vent moyen de composante OUEST sur le 60N revienne vers le 23 janvier 2019


Conclusion:

La date du pic de réchauffement est  devenue la référence pour le calcul du temps de réponse d'environ 14 jours pour le changement éventuel de la circulation dans la troposphère !

Dans ce cas-ci (deuxième réchauffement - RSS²) , le maximum du réchauffement a donc été atteint le 6 janvier 2019 !

Donc: 6/1/19 + +/-14jours = +/-20 janvier 2019

Une offensive hivernale commence à se dessiner de plus en plus après cette date dans les modèles globaux CFS, GFS et ECMWF !!!


Un réchauffement stratosphérique supplémentaire (RSS²) est apparu ce 5 janvier 2019 sur le nord-ouest de la Sibérie et est arrivé le 6 janvier à son maximum, avec -6° en son centre  .

Celui-ci forcera ensuite le vortex polaire principal qui , après s'être temporairement ressoudé sur l'Atlantique, à se splitter de nouveau vers le 11 janvier 2019 !

Vers le 13 janvier c'est une complète désorganisation de la circulation stratosphérique qui a lieu avec disparition du JETSTREAM STRATOSPHERIQUE POLAIRE !



PUBLIE PRECEDEMMENT

Suite au Réchauffement Stratosphérique Soudain (RSS) du 24/12/18 un déplacement, une déformation et deux splittings du Vortex polaire stratosphérique ont déjà eu lieu !


Evolution du réchauffement stratosphérique soudain (RSS) se produisant au Pôle nord vers 30km d'altitude 


Après avoir atteint son maximum, le 24 décembre 2018 vers 1h avec une température de +13° à 30 km d'altitude , le réchauffement stratosphérique sur le nord de la Sibérie a commencé à décliner lentement avec, ce 25 décembre à 10h, "seulement" +3° à la même altitude ; le vortex polaire était à ce moment-là encore centré sur l'Islande avec un minimum de -86° à la même altitude.

Ce 29 décembre le Vortex polaire a continué à complètement se déformer en s'étirant de plus en plus vers l'Amérique du nord.

Ce 31 décembre 2018 ce même vortex a éclaté comme prévu en 2 cellules dont la principale  s'étendant du sud du Groenland au sud de la Scandinavie et la secondaire sur le sud-ouest des USA.

Ce 1er janvier 2019, le courant moyen dans la haute stratosphère (10hPa) est passé de la composante OUEST (normale) vers la composante EST (anormale) .

Un autre splitting important (éclatement ) du Vortex polaire s'est manifesté le jeudi 3 janvier 2019 avec comme conséquence une complète désorganisation de la circulation stratosphérique avec une intensification des vents de composante EST dans la haute Stratosphère .

Observations concernant le 1er réchauffement stratosphérique soudain (RSS)

Dates à retenir concernant ce premier Réchauffement Stratosphérique Soudain (RSS)

en tenant compte du temps de réponse moyen (+/- 14 jours) pour l'impact potentiel du réchauffement stratosphérique soudain (RSS) et du/des splitting(s) entrainant des bouleversements des courants dans la stratosphère et ensuite dans la troposphère:

1. Le réchauffement stratosphérique maximum a eu lieu le 24/12/18 + 14jours =  7 janvier 2019



2. Le 1er split du vortex polaire dans haute stratosphère (10 hPa) a eu lieu le 31/12/18


3. Le début de l'inversion du vent zonal moyen (passage de la composante OUEST vers EST) sur le 60°N dans la haute stratosphère (10hPa) a eu lieu le 1/1/19


4. Le 2ième split du vortex polaire dans la haute stratosphère (10hPa) a eu lieu le 3/1/19 


5. Maximum du vent zonal moyen de composante EST sur le 60°N dans la haute stratosphère (10hPa) observé:

 le 7/1/19 (-9.7m/s)


Conclusion:

Comme  la date du pic de réchauffement serait la date de référence pour le calcul du temps de réponse d'environ 14 jours pour le changement éventuel de la circulation dans la troposphère !

Dans ce cas-ci (le premier réchauffement - RSS) , le maximum du réchauffement a été atteint le 24 décembre 2018!

Donc: 24/12/18 + 14jours = 7 janvier 2019

Après une première offensive hivernale bien neigeuse vécue depuis le début du mois de janvier 2019 (2/1/19) sur l'est et le centre de l'Europe, c'est l'ouest du continent qui va connaitre en milieu de semaine (9/1/19)  une petite offensive de quelques jours...



INFOS GENERALES:

Ce genre d'événement qui en soi n'a rien d'exceptionnel, est si intéressant car les éclatements ou les déplacements du vortex polaire ont régulièrement été suivis de périodes froides voire très froides dans l'hémisphère nord et donc également possibles sur l'Europe.

En effet, la perturbation et l'inversion des vents en stratosphère se communique généralement à la troposphère après environ 1 à 2 semaines.

Dès lors, notre habituel flux d'ouest caractéristique de nos hivers doux perd là un allié de poids dans le jetstream d'ouest situé dans la haute Troposphère en ralentissant.

Or, un flux d'ouest qui ralentit, ça laisse la place à d'autres flux dirigés vers nos régions, notamment des flux de nord ou d'est, acheminant de l'air beaucoup plus froid.

Quelques cas précédents:

*En décembre 1984, le vortex polaire a éclaté suite à un RSS.

Dans la haute stratosphère le vent à composante EST a remplacé, dès la fin du mois de décembre 1984, le vent zonal d'OUEST au 60N  .

C'est alors que l'on a relevé un vent moyen de -7m/s à 10hPa le 1er janvier 1985 avec un maximum atteint le 2 janvier 1985 avec une valeur moyenne de -14.1 m/s !

Ensuite ce vent à composante d'EST, s'étant propagé dans la Troposphère , a persisté longtemps provoquant ainsi un épisode hivernal rude sur nos régions.

Notes de l'IRM:

Du 1 au 4 janvier 1985, un hiver rigoureux s'est installé sur la Belgique.

Le 8, au-dessus des sables campinois, le minimum atteint -22,8°C à Kleine-Brogel (Peer) et, dans la vallée de la Lesse à Rochefort, le thermomètre descend à -25,4°C.

Le 14 janvier 1985, la température maximale ne dépasse pas -12,4°C à Thimister.

Le lendemain, elle atteint à peine -9,4°C à Middelkerke et -10,2°C à Stavelot.

Le 20 janvier 1985 : fin de l'une des vagues de froid les plus sévères que notre pays ait connues.

Entre le 5 et le 16 janvier 1985, on a compté huit jours avec des maxima inférieurs à -5°C à Uccle.

Cette période de grands froids fut marquée par un enneigement abondant sur tout le pays : on releva des épaisseurs de neige de 18 cm à Coxyde le 9, 23 cm à Uccle le 13, 24 cm à Kleine-Brogel (Peer) le 12 et jusqu'à 65 cm à Mont-Rigi (Waimes) le 13.

* En novembre 2010 le vortex polaire avait également  éclaté.

  Nous avons alors subi un mois de décembre 2010 froid et neigeux avec même un Noël blanc!

* Le 4 janvier 2013, le réchauffement stratosphérique était arrivé à son maximum avec +4° et le vortex polaire avait  éclaté le 8 janvier .

    Il s'en est suivi une deuxième partie d'hiver froide et régulièrement neigeuse qui s'est éternisée jusque début avril. A la mi-mars, nous connaissions même des chutes de neige très abondantes et des températures aussi basses que -15°C localement, ce qui est exceptionnel à ce moment-là de l'année.

*Le 4 février 2018, le réchauffement stratosphérique est à son maximum avec -12° à 30km d'altitude et le 9 février il y a éclatement du Vortex polaire avec une période très froide débutant vers le 19 février.

   Voici le communiqué de l'IRM à cet époque:

   "A partir du 16 février , l'anticyclone des Açores a commencé à se déplacer en direction du continent européen et a formé un début de jonction avec l'anticyclone situé au dessus du nord de la Scandinavie et de la Russie, avec pour conséquence de nous placer sous l'influence d'un temps plus froid mais lumineux causé par des masses d'air continentales froides.

   Le 19 février 2018, l'arrière d'une basse pression dont le centre était situé entre l'Islande et le Groenland s'est dirigé vers notre pays, mais la tendance vers un type de temps sec et froid s'est encore renforcée davantage.

   Du 21 février au 3 mars 2018, on peut décrire le temps de toute cette décade en Belgique comme très froid et  sec avec un ciel serein. Ceci a été provoqué par une puissante zone de haute pression au-dessus de la Scandinavie qui a envoyé vers notre pays des masses d'air continentales et ensuite polaires."

* Pour en revenir à la situation actuelle:

 Le modèle GFS semble déjà entrevoir ces modifications de flux pour la fin décembre 2018/début janvier 2019.                                                                                                                                                                                              Reste à voir si ce serait effectivement le cas, mais si nous étions amenés à connaître un hiver très froid, une partie des causes serait donc vraisemblablement à aller chercher dans la stratosphère.

REMARQUE:

Tous les éclatements ou déplacements de vortex polaire ne mènent cependant pas toujours  à une vague de froid chez nous, cela dépendra surtout du placement du ou des vortex et des anticyclones troposphériques.                                                                    

(Sources INFOMETEO- METEO BELGIQUE-METEOCIEL-IRM)


Publié le 18/12/18: 

Un Réchauffement Stratosphérique Soudain    (RSS) est attendu fin décembre            


Les dernières prévisions nous montrent de plus en plus la nouvelle possibilité d'un SSW (Sudden Stratosphéric Warming) ou d'un réchauffement stratosphérique soudain fin décembre 2018 et début janvier 2019 !

Ce genre de phénomène qui se produit entre 10 et 50 km d'altitude dans la Stratosphère, provoque le déplacement voire l'éclatement du VORTEX POLAIRE FROID (température <-80°c) en deux ou plusieurs parties qui sont repoussées par l'air chaud vers des latitudes moins élevées c'est à dire vers les nôtres.

Si l'une de ces poches froides vient s'installer au-dessus du continent européen, il est fort probable qu'une offensive hivernale digne de ce nom se déclenche environ 2 semaines après l'évènement stratosphérique.

En effet ce phénomène d'éclatement ou de déplacement perturbe fort le courant très rapide qui circule normalement d'ouest en est dans la Stratosphère , entre le début de l'automne et le début du printemps .

Comme le vortex polaire a la faculté d'impacter sur la direction des vents dans la Troposphère, ceux-ci seront également nettement perturbés suite au phénomène provoqué par le réchauffement stratosphérique.


A lire également:

The Influence of Stratospheric Vortex Displacements and Splits on Surface Climate

Daniel M. Mitchell, Lesley J. Gray, and James AnsteyNational Centre for Atmospheric Science, University of Oxford, Oxford, United Kingdom

April 2013


A strong link exists between stratospheric variability and anomalous weather patterns at the earth's surface. Specifically, during extreme variability of the Arctic polar vortex termed a "weak vortex event," anomalies can descend from the upper stratosphere to the surface on time scales of weeks.

Subsequently the outbreak of cold-air events have been noted in high northern latitudes, as well as a quadrupole pattern in surface temperature over the Atlantic and western European sectors, but it is currently not understood why certain events descend to the surface while others do not.

This study compares a new classification technique of weak vortex events, based on the distribution of potential vorticity, with that of an existing technique and demonstrates that the subdivision of such events into vortex displacements and vortex splits has important implications for tropospheric weather patterns on weekly to monthly time scales.

Using reanalysis data it is found that vortex splitting events are correlated with surface weather and lead to positive temperature anomalies over eastern North America of more than 1.5 K, and negative anomalies over Eurasia of up to −3 K. Associated with this is an increase in high-latitude blocking in both the Atlantic and Pacific sectors and a decrease in European blocking.

The corresponding signals are weaker during displacement events, although ultimately they are shown to be related to cold-air outbreaks over North America. Because of the importance of stratosphere-troposphere coupling for seasonal climate predictability, identifying the type of stratospheric variability in order to capture the correct surface response will be necessary.

Introduction

Polar stratospheric variability is largely dominated by vertically propagating Rossby waves of tropospheric origin (Andrews et al. 1987).

The resulting stratospheric anomalies can in turn descend and influence surface climate; although the mechanism for this coupling has received much attention over the past decade, no clear consensus has emerged.

Among the leading theories are wave-mean flow interactions (Christiansen 2001; Wittman et al. 2007) and wave reflections at the stratopause (Perlwitz and Harnik 2003).

The coupling time scales are, however, more accurately constrained, with northern annular mode (NAM) anomalies from weak and strong vortex events descending from the midstratosphere to the surface on time scales of weeks (Baldwin and Dunkerton 2001).

The resulting influence at the surface can cause midlatitude storms to become more intense, storm tracks to shift latitudinally, and the frequency of high-latitude blocking events to change (Thompson and Wallace 2001).

Traditionally, weak vortex events have been defined as either a "major sudden stratospheric warming," where a substantial fraction of the vortex air mass is rigorously mixed into the background flow, or a "minor sudden stratospheric warming," where the vortex air mass is disturbed but not to the same extent. Both types of event have often been defined using diagnostics based on the zonal mean (Charlton and Polvani 2007, hereafter CP07) or annular mode (Thompson and Wallace 1998).

However, recent research has shown increased understanding of these events when vortex centric diagnostics are used, such as 2D vortex moments (Waugh 1997; Waugh and Randel 1999; Mitchell et al. 2011a,b), which inherently take into account the zonally asymmetric nature of weak vortex events.

During weak vortex events the vortex can either be displaced off the pole (vortex displacement events) or split into two daughter vortices (vortex splitting events), and these are known to be predominantly associated with vertically propagating Rossby waves of wavenumber 1 and 2, respectively (Andrews et al. 1987).

The structure and evolution of the vortex during these types of events also differ greatly (Matthewman et al. 2009; Mitchell et al. 2011a) and may play an important role for understanding surface climate (Nakagawa and Yamazaki 2006).

We know from Hoskins et al. (1985) and Ambaum and Hoskins (2002) that a positive potential vorticity (PV) anomaly in the stratosphere will result in an elevated tropopause and vice versa.

Indeed a point change in stratospheric PV, Δq, can be linked to changes in the tropopause pressure, Δptrop, via the following relationship:

where Bu is the Burger number. We also note thatwhere f is the Coriolis parameter, is the relative vorticity, H refers to the horizontal component, and σ is a stratification-related mass density.

In the context of this study, high positive PV over the pole (i.e., the polar vortex) moving equatorward where there is lower ambient PV will result in a large positive PV anomaly in this region, which as a fractional change will be larger than the negative PV anomaly over the pole (i.e., where the vortex used to be).

This will be reflected in either a larger positive anomaly in or a bigger reduction in σ, or probably both (Ambaum and Hoskins 2002).

This movement of PV will broadly result in the following two features:

  1. a sinking of the tropopause over the pole where the vortex used to be, and

  2. an elevation of the tropopause at lower latitudes to where the vortex has been shifted.

As the PV anomaly is larger at lower latitudes, the change in tropopause height will also be greater in magnitude. However, the climatological tropopause height will be higher than in the polar region and hence communication to the surface may well be harder in this sense. Importantly, point 1 above will lead to similarities in surface influences for both splits and displacements, but point 2 will lead to differences due to the vortex residing at different latitudes and longitudes depending on the event type.

Wilcox et al. (2012) showed, using a blended thermal and dynamical tropopause definition, that significant variations in tropopause height were observed along the longitude plain as well as that of the latitude.

While they did not explicitly deal with the polar vortex, long-term trends in the tropopause height may well be associated with extreme vortex events.

If the tropopause is elevated in a certain region and depressed in another, then the change in thermal expansion and contraction of the troposphere may well influence surface climate.

With the possibility that weak vortex events will increase under climate change (Bell et al. 2009), as well as the possibility that the ratio of displacement to splitting events may also increase (Mitchell et al. 2012a,b), the need to understand surface influences for each type of event separately is becoming ever more crucial.

2. Methodsa.

Classification of event type

We use the 40-yr European Centre for Medium-Range Weather Forecasts (ECMWF) Re-Analysis (ERA-40) dataset over the period December 1958-April 2002. The data are available at 6-h time intervals and have 23 vertical pressure levels that range between 1000 and 1 hPa, with 12 of these levels representing the stratosphere.

Note that the analysis was also undertaken using National Centers for Environmental Prediction (NCEP)-National Center for Atmospheric Research (NCAR) data, and similar results were obtained.

Using ERA-40, we calculate elliptical diagnostics (Waugh 1997; Waugh and Randel 1999) of the Arctic polar vortex to obtain time series of the vortex area, aspect ratio, and centroid latitude on the 850-K isentropic surface (~10 hPa) (Mitchell et al. 2011a).

The results were insensitive to the choice of level between 650 and 1050 K (~30-5 hPa). This calculation involves identifying the PV contour that represents the vortex edge as the sharpest potential vorticity gradient in an equivalent latitude frame (Nash et al. 1996) and then applying PV weighting functions (2D vortex moments) inside the vortex region, defined in Cartesian coordinates as

where q is the potential vorticity, a is the moment order in the x direction, b is the moment order in the y direction (both a and b are nonnegative integers), and S is the surface of the vortex (Mitchell et al. 2011a).

To change from Cartesian to polar coordinate systems we use a polar stereographic projection (Waugh 1997),

where λ is the longitude and φ is the latitude.


Equation (3) allows for the calculation of time series of the following:

  1. The vortex area A, which is given by the zeroth-order moment, a = 0 and b = 0 (i.e., Aμ00).

  2. The vortex centroid latitude φcent, which is given by the first-order moment and is defined in Cartesian coordinates as . Transforming back to polar coordinates then yields

  3. The vortex aspect ratio r, which is given by the second-order moment such that

    where J denotes a transformation of Eq. (3) relative to the centroid of the vortex [see Matthewman et al. (2009) for more details].

Deseasonalized time series of A, φcent, and r were then run through a hierarchical clustering algorithm (Wilks 1995), which was able to correctly identify days in which the vortex was displaced, split, or stable, following the exact methodology of Hannachi et al. (2011) (the reader is referred to this study for technical details regarding the clustering algorithm).

To add confidence to the definition, we officially define an event as either a split or displacement if the vortex remains in this state for at least five consecutive days.1 If the vortex state changes between split and displaced (i.e., does not return to the stable state) within this 5-day window, a mixed event is defined.

A list of these events is given in Table 1, along with a comparison of weak vortex events defined in CP07.

Note that although we use the clustering algorithm to be consistent with Hannachi et al. (2011), similar dates can be achieved by using a simple threshold method in that splits are defined when the vortex aspect ratio is notably elliptical, and displacements when the centroid latitude is notably equatorward, adding confidence that the clustering algorithm is reliable in this case.b. Calculating the NAM

The NAM (known as the Arctic oscillation at the surface) is the leading mode of wintertime variability in the Northern Hemisphere circulation (Thompson and Wallace 1998; Baldwin 2001).

Here, we calculate the NAM as the leading empirical orthogonal function (EOF) of daily wintertime (November-April) geopotential anomalies poleward of 20°N.

The anomalies are calculated by subtracting the seasonal cycle, which has been smoothed with a 90-day low-pass filter.

The daily NAM anomalies are then determined by projecting daily geopotential anomalies onto the leading EOF patterns.

Finally, the NAM is normalized at each level so that the entire time series has unit variance. For this NAM definition we use the zonal-mean geopotential following Baldwin and Thompson (2009).

c. Calculating blocking

The blocking index is derived from daily-mean 500-hPa geopotential height Z500 according to the method of Tibaldi and Molteni (1990) as generalized to vary in both latitude and longitude by Scherrer et al. (2006).

At each horizontal grid point with latitude φ and longitude λ, the equatorward meridional gradient of Z500 is estimated as Δeqw = [Z500(λ, φ) − Z500(λ, φ − Δφ)]/Δφ, where Δφ = 15° as in Scherrer et al. (2006).

The poleward gradient is similarly defined as Δplw = [Z500(λ, φ + Δφ) − Z500(λ, φ)]/Δφ. An "instantaneous blocking" (IB) event is defined to occur when the following two conditions are fulfilled: 1) Δeqw > 0, indicating reversal of the climatological gradient of Z500 with easterlies equatorward of φ, and 2) Δplw < −10 m (° lat)−1, indicating strong westerlies poleward of φ.

The IB index b is defined to be 1 when these two conditions are satisfied, and 0 otherwise. The blocking frequency in events per day is simply the time mean of b.

d. Statistics

We use Student's t tests and Monte Carlo resampling methods to assess statistical significance throughout this study. The null hypothesis of the Student's t test is that the means of the datasets are not significantly different from zero. The validity of the test is also assessed by testing for Gaussianity of the data.

The Monte Carlo method assesses significance by comparing the probability density function (PDF) of the average NAM during 18 randomly averaged samples with those specifically of splitting and displacement events, resampled 105 times.

Here a sample is defined as a random period of 45 consecutive days during any winter in the dataset.

Note that 18 random events were used here because it is close to the composite of splitting events (18) and displacement events (19) determined during the classification stage. In reality the shape of the PDF does not vary greatly if 19 samples are used.

The result is also not sensitive to the number of resampling iterations.

3. Comparison with CP07

In this study, we distinguish between splitting and displacement vortex events using a clustering algorithm of the 850-K (~10 hPa) vortex area, aspect ratio, and centroid latitude (Mitchell et al. 2011a; Hannachi et al. 2011) to define 19 displacement events and 18 splitting events (see section 2).

Our method is completely distinct from that of CP07, who class an event as disturbed when the zonal mean zonal wind (ZMZW) at 60°N and 10 hPa reverses, and then proceed to class a split vortex when two vortices with a circulation ratio of 2:1 or higher are present (all other events are automatically defined as displacements).

Table 1 gives a comparison between dates defined using our method (column 2) and those defined in CP07 (column 5). The average cap (50°-90°N) temperature anomaly at 10 hPa for 5 days on either side of the event onset is also included to give a measure of event magnitude.

First, it is noted from Table 1 that our method identifies an expanded sample size of events relative to CP07. While many of the event dates are similar between the two studies, we note that the extra events included in this study are of high magnitude and seem to be as extreme as the previously identified events.

Figure 1 shows the cap temperature anomaly at 10 hPa for all events in Table 1 where our method defined an event, but CP07 did not (shown in blue), and all events in Table 1 where CP07 defined an event, but we did not (shown in red). The thick lines give composites of these events and on average our events are higher in magnitude than those of CP07, although we note that the differences are not statistically significant.

It should be noted that two events using our definition actually have a colder-than-average polar cap temperature (Fig. 1, solid black lines), and these occur simply because the vortex is particularly disturbed after the peak in polar cap temperature; in one case this eventually leads to another large polar cap temperature anomaly.

A further interesting disparity between the two sets of dates is the inclusion of six new events during the 1990s.

While traditionally thought of as a less dynamically active period than normal, this does not mean that extreme events were unable to occur. In particular the event on 12 January 1992 (event 40 in Table 1) was one of the largest on record in terms of cap temperature anomaly, reaching 12.8 K. Considering this event more closely,

Fig. 2 shows the evolution of the 850-K (~10 hPa) Northern Hemisphere PV fields at 2-day intervals following the onset. For reference a composite of PV during January is given in the top left panel.

When contrasted with the composite, it is clear that the polar vortex during this particular event is highly displaced from the pole.

Throughout the 8 days studied here the classic comma shape of the vortex can be observed, with a large filament of PV rotating out of the main vortex mass and being mixed into the background flow in a period of irreversible wave breaking.

Large displacements such as this are particularly important as they represent a shift in spatial location of high-magnitude stratospheric PV, which may influence the tropopause and hence tropospheric circulation (Ambaum and Hoskins 2002).

This particular result suggests that extreme vortex interactions over the 1990s may well be more abundant than previously thought.

Finally we compare the event type between our study and CP07.

Figure 3 shows the event type as a function of month and year for (top) our study and (bottom) CP07. We can see that in our study the majority of displacement events occur in February and March, whereas the majority of splits occur in December and January.

In contrast, the events from CP07 seem to be more evenly distributed throughout the year, although as in our definition there is a tendency for splits to be more concentrated in January.

To summarize we note the following advantages and disadvantages of using the method developed here compared with CP07:

Advantages:

  1. Our method uses a specifically developed vortex centric criterion in defining an event, and therefore takes into account the zonal asymmetry of the vortex evolution, whereas CP07 uses a method based on zonal symmetry.

  2. Taking measurements at a single point, as in CP07 (i.e., at 60°N and 10 hPa) means that events are often missed that can occur elsewhere (e.g., 65°N and 7 hPa). Our method makes use of full longitude-latitude fields, although we note that as in CP07 it does not have a vertical dependence.

  3. Using our definition often captures high-magnitude "minor" warmings, which can be more dynamically significant than "major" warmings defined in CP07. This allows for a higher sample size and therefore better statistics.

  4. Our method explicitly defines splits and displacements, whereas CP07 only defines splits and infers that all other events where the ZMZW at 60°N and 10 hPa is less than zero are displacements.

Disadvantages:

  1. The use of zonal-mean wind in CP07 means that multimodel comparisons can be made with little effort. On the other hand, few models output PV, as needed for our method, so more effort is required in first calculating this quantity.

  2. Our method depends on either a markedly elliptical vortex to define splits or extreme equatorward shifts in the vortex centroid to define displacements. If the vortex air mass becomes symmetrically disturbed about the pole, these diagnostics may well not define an event.

4. Analysis

Using our set of displacement and splitting events, the time-height evolution of the NAM over the winter period (October-March) is examined (Fig. 4).

Anomalies during a vortex splitting event (bottom) seem to have a greater influence on the surface than during a displacement event (top).

1 The signal during displacement events stops at the tropopause whereas it can descend to the surface during splitting events where it persists for ~60 days.

The vertical evolution between the two types of event also varies, with splitting events occurring almost instantaneously throughout the depth of the stratosphere, suggesting an excitation of the barotropic mode and lending support to the idea of wave resonance (Esler and Scott 2005).

The peak tropospheric signal occurs around 30 days following the event onset (we note that this result is not dominated by a few anomalous events) and suggests that if one has knowledge that a weak vortex event has begun, surface effects may be predictable on these time scales (Christiansen 2005).

However, a positive NAM anomaly is observed in the stratosphere as a precursor to both displacement and splitting events (Fig. 4), and is strong enough that an elevated tropopause is observed (Ambaum and Hoskins 2002) (solid black line).

Consequently, a further measure of predictability up to a month before the onset of these events may be apparent.

To determine the significance of the positive stratospheric NAM precursors, and subsequently its use for potential predictability, we test how likely it is that events of this magnitude occur over random periods during the winter.

This is assessed by randomly resampling the mean of the NAM over 45-day periods at 10 hPa during winter (Fig. 5a, PDF) and then comparing with the same measure calculated during splits (squares) and displacements (circles).

2  We observe that during both displacements and splits, the positive NAM signal is over two standard deviations from the mean of randomly sampled events and is therefore statistically significant at the 95% level.

Perhaps more important is the surface NAM [Arctic Oscillation (AO)] signal from each type of event and this is assessed using the same resampling technique, although now the surface signal for the period 15-60 days following (as opposed to preceding) an event is studied (Fig. 5b).

For splitting events (square) the AO has a mean value of −0.43, and is greater than two standard deviations from the mean of the PDF.

In contrast, the mean AO following displacement events (circle) is not significantly different from the mean of the PDF.

This implies that surface variability associated with the AO, such as an increased occurrence of high-latitude blocking and modulation of the midlatitude storms (Thompson and Wallace 2001), is far more likely following splitting events than displacements.

This result is in agreement with Nakagawa and Yamazaki (2006), who showed that events with enhanced upward flux of wavenumber-2 scale planetary waves were more likely to propagate to the surface than those with a reduced upward flux (see also Yoden et al. 1999).

However, it is noted that not all vortex splitting events are dominated by wave-2 activity.

It could be argued, however, that it is not only the instantaneous NAM response that matters when considering the tropospheric impact from a weak vortex event, but also the tendency in the NAM at the surface. For instance, one can see a positive AO anomaly during the displacement onset (Fig. 4a) followed by a negative AO anomaly ~30 days following the event.

Likewise for the splitting events, a weak negative AO anomaly is observed during the event onset that proceeds to become more strongly negative after ~30 days.

In terms of the AO trend these events would seem quite similar. We therefore plot the trend in the time series of the average AO for splits and displacements (Fig. 6, gray and black lines respectively).

3  The trends are calculated for ±15 days either side of each lagged day from the onset. For example, at lag = 0 we are calculating the trend in the AO for the period of 15 days before the event onset to 15 days after the event onset. At lag = 1 we do the same but for 14 days before the event onset to 16 days after, and so on.

Here we observe that both splits and displacements do have a similar AO trend around 10 days following the event onset, and this is significant according to Monte Carlo resampling at the 95% significance level (given by the dashed lines).

However, the AO values begin to recover earlier for displacements (at ~40 days after the onset) than for splits (~55 days), emphasizing the longer persistence time scales associated with splitting events.

To contrast the large-scale atmospheric dynamics between the two types of event and understand better the surface influences between them, we proceed to analyze the mean sea level pressure (MSLP) anomalies during three separate periods.

We choose the 30 days prior to the onset (precursor stage), 30 days after the onset (mature stage), and 30 days following the mature stage (decay stage) (Kolstad and Charlton-Perez 2010; Limpasuvan et al. 2004) (Fig. 7).

Note that the December-February (DJF) mean and variance fields in MSLP (Fig. 8) do not vary greatly between December and February, and therefore the DJF composites can be used with confidence to interpret winter anomalies.

The strongest MSLP anomaly is observed as a precursor to displacement events (Fig. 7a) and shows a wave-1 structure that projects well onto the stationary wave pattern (Garfinkel and Hartmann 2008), allowing for enhanced propagation of wave-1 anomalies into the stratosphere.

During the mature and decay phases following the displacement (Figs. 7d,g), very little surface signal is observed, consistent with the previous analysis.

A precursor signal is also observed before the splitting events (Fig. 7b), and while this does not project well onto the NAM, it does show wave-2 features and is consistent with previous studies (Garfinkel et al. 2010).

Over the two periods following the splitting event (Figs. 7e,h) an equatorward shift and a deepening is observed in both the Aleutian and Icelandic lows (for reference to the climatology, see Fig. 8), a pattern that is reminiscent of a negative NAM.

Consequently, it is likely that following splitting events storm tracks would shift equatorward and mobile cyclones would be enhanced at lower latitudes (Thompson and Wallace 2001).

It is also useful to consider the difference in the surface response for the two types of event relative to each other, rather than relative to the climatology.

Figure 7 (right) therefore plots the split minus displacement difference. Statistically the largest differences are observed in the precursor stage, implying that MSLP patterns are distinct preceding splits and displacements. However, for all 60 days following an event anomalously low pressures are observed over the northern Africa and western Europe regions, suggesting usefulness for a priori knowledge of an event type in seasonal forecasting over this region.

It should be noted here that CP07 did a similar analysis using their definitions of splits and displacements and found that a negative NAM response was present in both the splitting and displacement cases, albeit weaker than the negative NAM response reported in this study for splitting events.

No doubt the differences here arise in how we characterize events; however, more strong negative NAM events (or likewise fewer weak NAM events) make up the split composite in our study than do either of the composites in CP07.

To expand on this we consider regional projections of the surface temperature that are important for seasonal-scale climate forecasts (Fig. 9).

Consistent with the strong cyclonic flow of air around the anomalous low over North America, observed in the MSLP precursor to displacement events, temperature anomalies exceed 3 K over mainland North America (Fig. 9a).

The subsequent mature and decay phases show little significance except for high-magnitude cold anomalies of ~−1.5 K in the mid to high latitudes over the Americas (Fig. 9d).

Previous studies have noted the occurrence of "cold-air outbreaks" in this region (Thompson et al. 2002; Kolstad and Charlton-Perez 2010); however, this analysis shows that such outbreaks often occur following displacement events rather than splitting events. Because of the strong precursors associated with these events, a measure of predictability can be inferred.

In comparison, during the precursor stage of splitting events (Fig. 9b) cold anomalies are observed over North America and throughout Eurasia, whereas warmer anomalies are observed over Greenland.

The mature stage shows a different response to that of displacement events and indicates a −1-K anomaly over southern Eurasia.

However, the largest impact from either type of weak vortex event is evident during the decay period of a splitting event (Fig. 9h), during which a strong temperature dipole is observed with warm anomalies of up to 1.5 K over eastern North America, and cold anomalies of up to −3 K over northern Eurasia.

These strong negative anomalies are twice as large as the cold-air outbreaks noted over North America during displacement events and are unique to splitting events.

It is noted, however, that the surface temperature patterns are similar for both types of event during the decay phase, and it is the intensity of the signal that is most dissimilar.

As before, we also consider the difference (split minus displacement) in surface temperatures (Figs. 9c,f,i).

Consistent with the MSLP analysis, a strong difference is observed in the precursor stage, as well as a cold bias following a split compared to a displacement over North Africa and western Europe for all 60 days after the event, highlighting the influence of the stratospheric state on this region.

With a change in temperatures over lands and ocean an inevitable change in the land-sea contrast is observed, and hence Rossby wave generation can be modified.

This could potentially lead to a change in tropospheric blocking events (Andrews et al. 1987), which are known to result in persistently anomalous weather conditions.

To tie in the blocking with the NAM response we observed in Fig. 4, we first composite blocking activity for strong AO (AO > 1.6) and weak AO (AO < −1.6) events (Fig. 10; for a description of how we calculated the blocking index, see section 2c).

Here we see that in both cases the largest response is over the Atlantic and European regions.

Specifically, for the weak AO events (which are important for our analysis) increased blocking activity is observed over the North Atlantic, and decreased activity is observed in a band spanning from the mid-Atlantic to western Europe.

Considering the optimal periods where splits and displacements interact with the surface (i.e., from Fig. 4), we composite the instantaneous blocking for displacements, splits, and the difference (split minus displaced) in Fig. 11.

During the period before a displacement event we observe an increased occurrence of Eurasian blocking and decreased occurrence of blocking over the Atlantic and Pacific basins.

This is in agreement with Woollings et al. (2010), who show a similar spatial pattern to that of a positive AO (i.e., Fig. 10b), demonstrating good agreement with the positive NAM anomaly observed to descend from the stratosphere as a precursor to displacements.

Note that if a 5-day persistence criterion is imposed on the blocking definition, the blue region becomes less significant. Interestingly, the blocking activity following a displacement shows an increase over Canada that may well be linked to the cold-air outbreaks in this region (note that the blocking index that we are using shows Canada to be an area of low blocking activity).

While at the 95% level4 the significance of blocking activity is low for the period before splits (Fig. 11b), at the 90% level a significant increase is observed over the North Atlantic and northern Eurasia (not shown), hinting at a wave-2 type pattern.

The period following a split does, however, show more of a signal.

The spatial pattern suggests a strong negative AO (Fig. 10a) and agrees well with the downward propagation of a weak NAM signal observed during a vortex splitting event. In particular, large decreases in blocking activity can be seen over the Atlantic and European sectors and the blocking pattern over northern Eurasia is consistent with the large cold anomaly observed in Fig. 9.

The split minus displacement difference (Fig. 11f) in this region is also large and, when taken in conjunction with the surface temperature and MSLP analyses, it is clear that this is a region of importance when considering the different influences of weak vortex events on surface climate.

5. Summary

In this paper we have developed a novel method of defining polar vortex splits and displacement.

This method has been contrasted against that of Charlton and Polvani (2007) to reveal advantages and disadvantages of using both, and most importantly we have shown that one must treat vortex splitting and displacement events individually if a true representation of the subsequent surface influence is to be achieved.

To do this the most up-to-date measures of vortex variability and tropospheric blocking have been employed and yield the following conclusions:

a. Vortex displacements
  • Preceding these events are often anomalously low pressure systems over North America and high pressure systems over western Europe and the Pacific. Associated with this are warm temperature anomalies over Northern America and an increase in blocking over northern Eurasia.

  • While the stratospheric NAM anomaly is large for these events, with a potential for predictability up to a month before hand, the anomaly is not seen to descend through the troposphere. At the surface the AO trend is similar around the onset date for both splits and displacements, although the AO anomaly persists for ~15 days less during displacements.

  • The largest surface impact from displacement is observed over the month following an event and shows anomalously cold temperatures of magnitude −1.5 K over North America, a feature that is not observed for the splitting case. Associated with this, increased blocking activity is observed over Canada.

b. Vortex splits
  • Preceding splitting events are anomalously low temperatures over Eurasia, with a wave-2-like pattern observed in MSLP.

  • The midstratospheric NAM signal following splitting events is weaker than that which follows displacements events, but importantly anomalies can descend from the midstratosphere to the surface, unlike displacement events. The evolution of the anomalies are also far more barotropic than during displacement events.

  • For 60 days following a splitting event a coherent negative AO anomaly is observed. Consistent with this, high-latitude blocking in both the Atlantic and Pacific basins increases while blocking in the mid-Atlantic, Europe, and western Eurasia decreases. Ultimately the largest effect from these events is observed over northern Eurasia with low temperature anomalies of up to −3 K.

Recently many studies have alluded to the stratospheric involvement in extended range forecasting (e.g., Christiansen 2005; Fletcher et al. 2007; Hardiman et al. 2011).

The implications of these results for monthly-scale climate forecasts in the high northern latitudes are great and the different surface response to displaced and split vortex events demonstrates the necessity for forecasting systems, and climate models, to correctly simulate the evolution and frequency of these two types of vortex disturbances.

On a fundamental level this will involve models including a fully resolved stratosphere with an excellent representation of how the structure and evolution of the Arctic polar vortex varies throughout winter, so that the distinct influence from splitting and displacement events can be appropriately captured.

Acknowledgments

We are grateful for the insightful comments from Chaim Garfinkel and the two anonymous reviewers.

We would also like to thank Brian Hoskins and Lorenzo Polvani for discussions regarding this study.

DMM and JA are supported by a grant from the UK Natural Environmental Research Council (NERC) and LJG is funded by the NERC National Centre for Atmospheric Science (NCAS).

REFERENCES Ambaum, M., and B. Hoskins, 2002: The NAO troposphere-stratosphere connection. J. Climate, 15, 1969-1978. Link, Google Scholar
Andrews, D., J. Holton, and C. Leovy, 1987: Middle Atmosphere Dynamics. Academic Press, 489 pp. Google Scholar
Baldwin, M., 2001: Annular modes in global daily surface pressure. Geophys. Res. Lett., 28, 4115-4118. Crossref, Google Scholar
Baldwin, M., and T. Dunkerton, 2001: Stratospheric harbingers of anomalous weather regimes. Science, 294, 581-584. Crossref, Google Scholar
Baldwin, M., and D. Thompson, 2009: A critical comparison of stratosphere-troposphere coupling indices. Quart. J. Roy. Meteor. Soc., 135, 1661-1672. Crossref, Google Scholar
Bell, C., L. Gray, A. Charlton-Perez, M. Joshi, and A. Scaife, 2009: Stratospheric communication of El Niño teleconnections to European winter. J. Climate, 22, 4083-4096. Link, Google Scholar
Charlton, A., and L. Polvani, 2007: A new look at stratospheric sudden warmings. Part I: Climatology and modeling benchmarks. J. Climate, 20, 449-469. Link, Google Scholar
Christiansen, B., 2001: Downward propagation of zonal mean zonal wind anomalies from the stratosphere to the troposphere-Model and reanalysis. J. Geophys. Res., 106, 27 307-27 322. Crossref, Google Scholar
Christiansen, B., 2005: Downward propagation and statistical forecast of the near-surface weather. J. Geophys. Res.,110, D14104, doi:10.1029/2004JD005431. Google Scholar
Esler, J., and R. Scott, 2005: Excitation of transient Rossby waves on the stratospheric polar vortex and the barotropic sudden warming. J. Atmos. Sci., 62, 3661-3682. Link, Google Scholar
Fletcher, C., P. Kushner, and J. Cohen, 2007: Stratospheric control of the extratropical circulation response to surface forcing. Geophys. Res. Lett.,34, L21802, doi:10.1029/2007GL031626. Google Scholar
Garfinkel, C., and D. Hartmann, 2008: Different ENSO teleconnections and their effects on the stratospheric polar vortex. J. Geophys. Res., 113, D18114, doi:10.1029/2008JD009920. Crossref, Google Scholar
Garfinkel, C., D. Hartmann, and F. Sassi, 2010: Tropospheric precursors of anomalous Northern Hemisphere stratospheric polar vortices. J. Climate, 23, 3282-3299. Link, Google Scholar
Hannachi, A., D. Mitchell, L. Gray, and A. Charlton-Perez, 2011: On the use of geometric moments to examine the continuum of sudden stratospheric warmings. J. Atmos. Sci., 68, 657-674. Link, Google Scholar
Hardiman, S. C., and Coauthors, 2011: Improved predictability of the troposphere using stratospheric final warmings. J. Geophys. Res., 116, D18113, doi:10.1029/2011JD015914. Crossref, Google Scholar
Hoskins, B., M. McIntyre, and A. Robertson, 1985: On the use and significance of isentropic potential vorticity maps. Quart. J. Roy. Meteor. Soc., 111, 877-946. Crossref, Google Scholar
Kolstad, E., and A. Charlton-Perez, 2010: Observed and simulated precursors of stratospheric polar vortex anomalies in the Northern Hemisphere. Climate Dyn., 37, 1443-1456. Google Scholar
Limpasuvan, V., D. Thompson, and D. Hartmann, 2004: The life cycle of the Northern Hemisphere sudden stratospheric warmings. J. Climate, 17, 2584-2596. Link, Google Scholar
Matthewman, N., J. Esler, A. Charlton-Perez, and L. Polvani, 2009: A new look at stratospheric sudden warmings. Part III: Polar vortex evolution and vertical structure. J. Climate, 22, 1566-1585. Link, Google Scholar
Mitchell, D., A. Charlton-Perez, and L. Gray, 2011a: Characterizing the variability and extremes of the stratospheric polar vortices using 2D moment analysis. J. Atmos. Sci., 68, 1194-1213. Link, Google Scholar
Mitchell, D., L. Gray, and A. Charlton-Perez, 2011b: The structure and evolution of the stratospheric vortex in response to natural forcings. J. Geophys. Res.,116, D15110, doi:10.1029/2011JD015788. Google Scholar
Mitchell, D., S. Osprey, L. Gray, N. Butchart, S. Hardiman, A. Charlton-Perez, and P. Watson, 2012a: The effect of climate change on the variability of the Northern Hemisphere stratospheric polar vortex. J. Atmos. Sci., 69, 2608-2618. Link, Google Scholar
Mitchell, D., and Coauthors, 2012b: The nature of arctic polar vortices in chemistry-climate models. Quart. J. Roy. Meteor. Soc., 138, 1681-1691. Crossref, Google Scholar
Nakagawa, K., and K. Yamazaki, 2006: What kind of stratospheric sudden warming propagates to the troposphere? Geophys. Res. Lett.,33, L04801, doi:10.1029/2005GL024784. Google Scholar
Nash, E., P. Newman, J. Rosenfield, and M. Schoeberl, 1996: An objective determination of the polar vortex using Ertel potential vorticity. J. Geophys. Res., 101, 9471-9478. Crossref, Google Scholar
Perlwitz, J., and N. Harnik, 2003: Observational evidence of a stratospheric influence on the troposphere by planetary wave reflection. J. Climate, 16, 3011-3026. Link, Google Scholar
Scherrer, S., M. Croci-Maspoli, C. Schwierz, and C. Appenzeller, 2006: Two-dimensional indices of atmospheric blocking and their statistical relationship with winter climate patterns in the Euro-Atlantic region. Int. J. Climatol., 26, 233-249. Crossref, Google Scholar Thompson, D., and J. Wallace, 1998: The Arctic Oscillation signature in the wintertime geopotential height and temperature fields. Geophys. Res. Lett., 25, 1297-1300. Crossref, Google Scholar
Thompson, D., and J. Wallace, 2001: Regional climate impacts of the Northern Hemisphere annular mode. Science, 293, 85-89. Crossref, Google Scholar
Thompson, D., M. Baldwin, and J. Wallace, 2002: Stratospheric connections to Northern Hemisphere wintertime weather: Implications for prediction. J. Climate, 15, 1421-1428. Link, Google Scholar
Tibaldi, S., and F. Molteni, 1990: On the operational predictability of blocking. Tellus, 42A, 343-365. Crossref, Google Scholar
Waugh, D., 1997: Elliptical diagnostics of stratospheric polar vortices. Quart. J. Roy. Meteor. Soc., 123, 1725-1748. Crossref, Google Scholar
Waugh, D., and W. Randel, 1999: Climatology of Arctic and Antarctic polar vortices using elliptical diagnostics. J. Atmos. Sci., 56, 1594-1613. Link, Google Scholar
Wilcox, L., B. Hoskins, and K. Shine, 2012: A global blended tropopause based on ERA data. Part I: Climatology. Quart. J. Roy. Meteor. Soc., 138, 561-575. Crossref, Google Scholar
Wilks, D., 1995: Statistical Methods in the Atmospheric Sciences: An Introduction. Academic Press, 467 pp. Google Scholar
Wittman, M., A. Charlton, and L. Polvani, 2007: The effect of lower stratospheric shear on baroclinic instability. J. Atmos. Sci., 64, 479-496. Link, Google Scholar
Woollings, T., A. Charlton-Perez, S. Ineson, A. Marshall, and G. Masato, 2010: Associations between stratospheric variability and tropospheric blocking. J. Geophys. Res.,115, D06108, doi:10.1029/2009JD012742. Google Scholar
Yoden, S., T. Yamaga, S. Pawson, and U. Langematz, 1999: A composite analysis of the stratospheric sudden warmings simulated in a perpetual January integration of the Berlin TSM GCM. J. Meteor. Soc. Japan, 77, 431-445. Google Scholar

1 Applying persistence time scales of between 5 and 10 days does not alter the conclusions of this study.

2 This period was chosen to allow time for anomalies to propagate to the surface following an event onset. Changing the period decreases the magnitude of the signal, but crucially does not alter the significance.

3 The trend is calculated by applying a least squares linear fit to the time series.

4 We choose to use a Monte Carlo method of significance testing (see section 2) because the underlying distribution of blocking activity is not assumed to be Gaussian.