Enhanced Seasonal Prediction of European Winter Warming Following Volcanic Eruptions

6168 JOURNAL OF CLIMATE VOLUME 22

Enhanced Seasonal Prediction of European Winter Warming following Volcanic Eruptions

A. G. MARSHALL,* A. A. SCAIFE, AND S. INESON Hadley Centre, Met Ofﬁce, Exeter, United Kingdom

(Manuscript received 18 March 2009, in ﬁnal form 14 July 2009)

ABSTRACT

The impact of explosive volcanic eruptions on the atmospheric circulation at high northern latitudes is assessed in two versions of the Met Ofﬁce Hadley Centre’s atmospheric climate model. The standard version of the model extends to an altitude of around 40 km, while the extended version has enhanced stratospheric resolution and reaches 85-km altitude. Seasonal hindcasts initialized on 1 December produce a strengthening of the winter polar vortex and anomalous warming over northern Europe characteristic of the positive phase of the Arctic Oscillation (AO) when forced with volcanic aerosol following the 1963 Mount Agung, 1982 El Chicho´ n, and 1991 Mount Pinatubo eruptions, as is observed. The AO signal in the extended model is of comparable strength to that in the standard model, showing that there is little impact from both increasing the vertical resolution in the stratosphere and extending the model domain to near the mesopause. The presence of this signal in the models, however, is likely due to the persistence of the observed signal from the initial conditions, because a similar set of experiments initiated with the same conditions, but with no volcanic aerosol forcing, exhibits a similar response as the forced runs. This suggests that the model has limited ﬁdelity in capturing the response to volcanic aerosols on its own, consistent with previous studies on the impact of volcanic forcing in long climate simulations, but does support the premise that seasonal winter forecasts are substantially improved with the inclusion of stratospheric information.

1. Introduction and Brownscombe 1983; Angell 1993) resulting from the absorption of terrestrial longwave and solar near- The three largest explosive volcanic eruptions to have infrared radiation (e.g., Newell 1970; Pollack et al. 1976; occurred in the second half of the twentieth century Kinne et al. 1992). Thermal fluctuations related to the were Mount Agung in Bali, Indonesia (March and May tropical quasi-biennial oscillation (QBO; e.g., Angell 1963), El Chicho´ n in Chiapas, Mexico (March–April and Korshover 1964) also contributed to the lower- 1982), and Mount Pinatubo in the Luzon, Philippines stratospheric warming; however, the total warming has (June 1991). These tropical eruptions emitted tens of been shown to considerably exceed that associated megatons total of sulfur dioxide (SO ) and particulate 2 with the QBO (e.g., Parker and Brownscombe 1983; material into the lower stratosphere (e.g., McInturff et al. Stenchikov et al. 1998). The SO molecules that were 1971; McCormick and Veiga 1992; Baran and Foot 1994; 2 injected into the stratosphere oxidized within a few Robock 2002) and caused a warming of the tropical lower weeks to form sulfate aerosols that had both heating and stratosphere on the order of a few degrees Celsius (e.g., cooling effects on the troposphere and surface below, Angell and Korshover 1978; Labitzke et al. 1983; Parker with increased emissions of downward longwave radiation (heating) and reduced downward visible and near- infrared fluxes (cooling). The total radiative balance was * Current affiliation: Centre for Australian Weather and Cli- mate Research, Hobart, Tasmania, Australia. affected by the distributions of clouds, water vapor, and surface temperature, which varied in response to the perturbed tropospheric climate (Stenchikov et al. 1998; Corresponding author address: Andrew G. Marshall, Centre for Ramachandran et al. 2000), but the net effect was a sub- Australian Weather and Climate Research, CSIRO Marine and Atmospheric Research, Castray Esplanade, Hobart, TAS 7000, tropical and tropical cooling at the surface for about 2 yr Australia. after the eruptions (Robock and Mao 1995); however, E-mail: [email protected] this cooling was not observed in the first year after the

DOI: 10.1175/2009JCLI3145.1

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1982 El Chichoń eruption because the developing 1982/ not account for the QBO, which is known to have a sig- 83 El Ninõ event produced a large compensating warm- nificant effect on the winter circulation in the Northern ing (Robock 2000). Hemisphere extratropics (e.g., Holton and Tan 1980, The eruptions’ impacts were global. The volcanic sulfate 1982; Dunkerton and Baldwin 1991). Ramachandran aerosol from the Mount Agung eruption spread mostly et al. (2000) used the Geophysical Fluid Dynamics Lab- throughout the Southern Hemisphere, the El Chichoń oratory (GFDL) ‘‘SKYHI’’ GCM with Mount Pina- aerosol spread mostly throughout the Northern Hemi- tubo aerosol parameters similar to those developed by sphere, and the Mount Pinatubo aerosol spread to both Stenchikov et al. (1998). They generated substantial poles (Viebrock and Flowers 1968; Stenchikov et al. 1998). longwave stratospheric cooling in the middle to high lat- Both the stratosphere and troposphere responded dy- itudes that contributed to the enhanced equator-to-pole namically to the perturbed radiative properties of the temperature gradient initiated by the tropical lower- tropical atmosphere following the eruptions. Climate stratospheric warming. The impact on the polar vortex variations included a substantial winter surface warm- was not assessed, however, because of the relatively high ing of Northern Hemisphere continents, including over interannual variability and low signal-to-noise ratio in Europe and Siberia and anomalous cooling over the high-latitude stratospheric temperature anomalies com- Middle East and Greenland, as was documented fol- pared to those observed. Further, their model also ne- lowing the El Chicho´ n and Mount Pinatubo eruptions glected the QBO. Neither the Stenchikov et al. (1998) (Robock and Mao 1992; Stenchikov et al. 1998, respec- nor Andronova et al. (1999) modeling studies generated tively). The most widely accepted large-scale dynamical strong extratropical longwave cooling following the mechanism associated with these anomalies stems from Mount Pinatubo eruption; Ramachandran et al. (2000) the lower-stratospheric warming in the tropics, which attributed this cooling effect to the impact of high clouds gives rise to an enhanced equator-to-pole temperature in increasing the thermal emission by the aerosol layer. gradient. This produces an enhanced northern polar Stenchikov et al. (2002) examined the sensitivity of vortex and a response in the tropospheric circulation that the GFDL SKYHI GCM to short-term ozone loss af- gives rise to large-scale surface climate variations (e.g., ter the Mount Pinatubo eruption, which is caused by Graf et al. 1993; Kodera 1994). This dynamical effect has heterogeneous and radiative volcanic aerosol effects been described as a modification of the interaction be- on ozone photochemistry (e.g., Hofmann and Solomon tween stratospheric westerlies and vertically propagating 1989; Solomon et al. 1998). Experiments that used post– planetary waves (Perlwitz and Graf 1995), in that an en- Mount Pinatubo ozone anomalies produced a strong hanced winter polar vortex leads to increased refraction of positive phase of the AO comparable to that observed, planetary waves, which in turn decreases the deceleration and a weaker positive AO response was produced by of the vortex (e.g., Stenchikov et al. 2006). Thus, the im- experiments that included only the tropospheric effects pact of volcanic aerosols leads to an enhanced positive of aerosols (i.e., no stratospheric heating); however, in- phase of the Arctic Oscillation (AO) and associated North terpretation of the results was again limited by the ab- Atlantic Oscillation (NAO; Thompson and Wallace 1998) sence of a QBO in the model. The importance of the that is most prominent in boreal winter and persists for up QBO was subsequently demonstrated by Stenchikov to 2 yr after each eruption (e.g., Robock and Mao 1992; et al. (2004) in simulations that included post–Mount Stenchikov et al. 2006). Pinatubo volcanic aerosols and a realistic QBO simu- The observed AO response to volcanic eruptions has lation, but without volcanically induced ozone deple- been reproduced with varying success in general circu- tion. The winter response to the combined effect of lation models (GCMs) over the last decade. Kirchner aerosols and the QBO was a positive AO phase that was et al. (1999) reproduced a Northern Hemisphere winter enhanced by the westerly QBO phase during the 1992/93 warming pattern in the troposphere of similar strength to winter, as observed. that observed using the ECHAM4 GCM with Mount More recently Jones et al. (2005) demonstrated the Pinatubo aerosol parameters developed by Stenchikov potential impact of a strengthened equator-to-pole et al. (1998). However, the intensity of the stratospheric stratospheric temperature gradient on the winter vortex polar vortex anomaly was much weaker than that ob- following a generic volcanic ‘‘supereruption’’ (on the served, and its position and strength were strongly influ- scale of 100 times larger than that of Mount Pinatubo), enced by sea surface temperature (SST) forcing. This the last of which was the Toba eruption in Sumatra, made it difficult to assess the effect of the stratosphere Indonesia (around 71,000–75,000 years ago; see Rose on the tropospheric circulation, and internal variability and Chesner 1990; Zielinski et al. 1996; Oppenheimer likely played a role in the five-member ensemble mean 2002). Tropical lower-stratospheric warming from vol- surface temperature response. Further, their model did canic aerosol prescribed in their experiment increased

Unauthenticated | Downloaded 09/30/21 01:51 PM UTC 6170 JOURNAL OF CLIMATE VOLUME 22 the meridional temperature gradient, which led to a in vegetation, sulfur, soot, and biomass emissions. Sea strong positive phase of the AO and subsequent winter surface temperature and sea ice extent variations are warming over northern landmasses in the first two win- specified from an analysis of historical observations ters following the eruption. The winter warming was at- (Rayner et al. 2003), and atmospheric ozone concen- tributed in equal parts to the positive AO mode (which trations are held constant at 1990 levels. dominated the response south of the Arctic Circle) and Three experiments are devised for the L38 and L60 increased longwave forcing from the stratospheric aero- models. The first of these incorporates stratospheric sols (which dominated north of the circle). The high- volcanic aerosol column mass into winter [December– latitude response to nineteenth- and twentieth-century February (DJF)] hindcasts for 1963/64, 1964/65, 1982/83, volcanic eruptions simulated in many present-day GCMs 1983/84, 1991/92, and 1992/93. As discussed, the main remains poor, however. Stenchikov et al. (2006) assessed atmospheric thermal and dynamical effects of recent the response to nine volcanic eruptions during the pe- volcanic eruptions have been shown to persist for about riod 1860–1999 in long climate simulations using seven 2 yr after each eruption, and thus we produce L38 and climate models included in the model intercomparison L60 model hindcasts for the two boreal winters following conducted as part of the Intergovernmental Panel on the three largest explosive volcanic eruptions that have Climate Change (IPCC) Fourth Assessment Report occurred over the last 50 yr: Mount Agung (March–May (AR4; Alley et al. 2007). The models tended to simulate 1963), El Chichoń (March–April 1982), and Mount Pina- a positive AO phase in response to volcanic forcing, how- tubo (June 1991). Fifteen-member ensembles are pro- ever this was much weaker than that observed. Stenchikov duced for each of these six ‘‘postvolcanic’’ winter periods. et al. (2006) noted the coarseness of the models’ reso- The atmospheric conditions for each ensemble are ini- lution and suggested that an improved response may be tialized using 6-hourly data from the 40-yr European achieved by both increasing the vertical resolution in Centre for Medium-Range Weather Forecasts (ECMWF) the stratosphere and extending the model lid into the Re-Analysis (ERA-40; Uppala et al. 2005), starting from mesosphere. 1200 UTC 27 November and ending at 0000 UTC This study uses two versions of the Met Office Hadley 1 December, to produce 15 hindcasts that are integrated Centre climate model with different vertical resolutions from a model start date of 1 December. above the tropopause to investigate the extratropical Stratospheric volcanic aerosol column mass totals are winter response to volcanic aerosol forcing and its sen- calculated from the updated monthly optical depth sitivity to stratospheric resolution in seasonal hindcasts. dataset of Sato et al. (1993; updates are documented Model experiments and observational data are described online at http://www.giss.nasa.gov/data/strataer/). The in section 2, results are presented in section 3, and con- aerosol is applied globally to the model across four lat- clusions are summarized in section 4. itude bands of approximately equal areas (908–308S, 308S–08,08–308N, and 308–908N) and above the tropopause as a simple step function with an equal mass 2. Data description mixing ratio (MMR) at all altitudes; thus, the aerosol mass density decreases with altitude. The ejected Mount a. Model experiments Pinatubo volcanic aerosol was confined to the lower- We make use of the 38- and 60-level atmosphere-only and midstratosphere with the bulk of the aerosol cloud configurations of the first-generation Hadley Centre below 30 mb (Stenchikov et al. 1998; Andronova et al. Global Environmental Model (HadGEM1; Martin 1999), and hence the aerosol is applied in our experi- et al. 2006), with many of the changes proposed for ments from the tropopause to 30 mb. The radiative ef- HadGEM2-A as documented in Collins et al. (2008), fect of aerosols is modeled by scattering and absorbing and a spatial resolution of 1.258 latitude 3 1.8758 lon- incoming solar radiation; most of the scattering is at the gitude. The 38-level version (L38) has a model top in the shorter wavelengths and most of the absorption is at the stratosphere at 39.3 km (;5 mb) and the 60-level ver- longer wavelengths including near-infrared. The aerosol sion (L60) has a model top near the mesopause at also absorbs longwave radiation emitted by the surface 84.1 km (;0.005 mb). Both models have the same ver- and troposphere. Factors such as volcanic ash, dust, tical resolution in the troposphere, a horizontal resolu- smoke, and changes in tropospheric aerosols and their tion of N96, and a time step of 20 min, allowing for cloud–aerosol interactions are not incorporated because a clean assessment of the impact of volcanic aerosol only the effects of changes in stratospheric aerosols forcing. All model experiments described here are forced are included (Jones et al. 2005). The time evolution of with time-varying greenhouse gas concentrations, in- aerosol distributed across the four latitude bands is cluding CO2,CH4,N2O, CFCl3, and CF2Cl2, and changes shown in Fig. 1 for each volcanic eruption.

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FIG. 1. Time evolution of volcanic stratospheric aerosol applied globally to the model across four latitude bands for each eruption. The ﬁrst major explosive eruption for each event occurs during month 2. Latitude bands have approximately equal area: 1) 308–908N, 2) 08–308N, 3) 308 S–08, and 4) 908–308S. Shading corresponds to stratospheric aerosol optical depth values at 0.55 mm of 0–0.05 (white), 0.05–0.10 (light gray), 0.10–0.15 (dark gray), and 0.15–0.20 (black).

The second experiment provides a ‘‘nonvolcanic’’ b. Observational data winter climatology for calculating extratropical anomalies in response to volcanic forcing. We thus produce We use the National Centers for Environmental 5-month model hindcasts for a subset of years during Prediction–National Center for Atmospheric Research which no volcanic eruptions occurred: 1968/69, 1987/88, (NCEP–NCAR) reanalysis 1 dataset (NNR1; Kalnay 1989/90, 1995/96, and 1997/98. Note that there is no net et al. 1996) for model comparison and assessment. AO signal averaged over these winters. These years are NNR1 is a global reanalysis spanning the period 1948– also chosen so that composite anomalies calculated 2005, produced with a model resolution of T62L28, and relative to climatology are not biased by cold/warm provides global, quality-controlled datasets using a El Nin˜ o–Southern Oscillation (ENSO) episodes, east- ‘‘frozen’’ data assimilation–forecast system. The use of erly/westerly QBO phases, or stratospheric sudden such a system prevents problems of pseudoclimate sig- warming events, which may otherwise contaminate the nals being introduced into the dataset through changes anomalous postvolcanic signal in the extratropics (as- in assimilation techniques and model formulation. suming a linear superposition) by modulating the north- To isolate the observed volcanic signal we chose non- ern winter polar vortex (e.g., Hamilton 1993; Baldwin volcanic winter reference periods between 1956 and et al. 2001). Fifteen-member ensembles for each of 1991 during which no volcanic eruptions occurred and these nonvolcanic winter periods are also initialized at the atmosphere was relatively clear of volcanic aerosol 6-hourly intervals starting from 1200 UTC 27 November [such that global mean monthly optical thicknesses at and ending with 0000 UTC 1 December to produce 550 nm were less than 0.015 (from Sato et al. 1993)]. hindcasts that are integrated from a model start date of These DJF periods fell within December 1957–February 1 December. 1961, December 1971–February 1974, December 1976– The third experiment, referred to as ‘‘NOVOLC,’’ is February 1981, and December 1985–February 1991. With a repeat of the ﬁrst experiment (i.e., for the six post- the exception of the earliest four winters, these periods volcanic winters), but without volcanic aerosol forcing overlap with the recent nonvolcanic reference periods in the models. This additional set of hindcasts allows us used by Stenchikov et al. (2006). Observed volcanic to separate the impact of volcanic aerosol from the role anomaly composites are thus calculated in a manner of initial/boundary condition forcing in the hindcasts similar to that used for the ensemble mean composites (section 3c). as postvolcanic minus nonvolcanic winter averages, and

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FIG. 2. Number of six-member composites (y) that pertain to each anomaly bin (x) for randomly sampling single model realizations, without the forced AO response to volcanic eruptions, over six winters. A total of 156 winter (DJF) composites of anomalous 50-mb geopotential height averaged poleward of 658N are sampled for L38 and L60. Values lower than the observed anomaly are shaded gray, and the number of composites (expressed as a percentage of the total) that reproduce or exceed the strength of the observed response to volcanic aerosol is shown in parentheses in the title for each plot. Every third tick mark on the x axis is labeled, with a bin width of 20 m. are also not biased by cold/warm ENSO episodes, to volcanic eruptions. We repeat this for all possible easterly/westerly QBO phases, or stratospheric sudden single-realization combinations to generate 156 winter warming events. Note that fewer years are used in the composites of anomalous geopotential height at 50 mb model climatology (5) than are used in the observed (Z50) for both the L38 and L60 models. This is quanti- climatology (18) because of the limitations imposed by ﬁed and presented in a meaningful way using a bar chart computational expense. However with 15 hindcasts per metric for which Z50 anomalies averaged poleward of ensemble, the total number of realizations in the model 658N are calculated from each composite and assigned climatology exceeds the total number of winters used in to frequency bins. The number of six winter samples the observed climatology and provides a high signal-to- (expressed as a percentage of the total) that reproduce noise ratio in our analyses. or exceed the strength of the observed response is used to assess the likelihood of the observed signal being due to internal variability, for which we adopt the parlance 3. Model response to volcanic aerosol forcing used by Alley et al. (2007): The strong observed AO signal is extremely unlikely to be an artifact of internal a. Robustness of observed signal variability (or extremely likely to be a robust response to The strength of the observed extratropical response to volcanic forcing) if it is reproduced in less than 5% of the volcanic forcing is not well represented in present-day model iterations. climate simulations despite the models’ ability to re- The results are presented in Fig. 2 for the L38 and L60 produce radiative warming of the lower tropical strato- models. Note that the integrated Z50 anomaly over all sphere (e.g., Stenchikov et al. 2006). This suggests bins gives an average polar Z50 anomaly near zero. The possible model deﬁciencies in capturing the subsequent percentage of iterations that reproduce or exceed the dynamical response that affects high-latitude circulation strength of the observed response for each metric are anomalies; however, there may be an alternative ex- shown in parentheses in the title for each plot. The ob- planation. It is possible that the observed signal is an served strength of winter Z50 anomalies is reproduced artifact of internal variability rather than a volcanic in only 0.9% and 2.2% of iterations for L38 and L60, forced signal. If this were the case, it would be possible respectively, and it is therefore extremely unlikely that to readily reproduce the observed lower-stratospheric the observed signal is due to internal variability. We AO signal in a model without volcanic forcing using further show that HadGEM1 reproduces realistic AO a single realization for each of the six winters. We test variability by conducting a covariance EOF analysis this here by randomly sampling single-model realiza- of Z50 for experiments without volcanic forcing; the tions over the six winters to produce composites of six leading EOF for the models is similar to that for NNR1, realizations from which we exclude the forced response with a peak magnitude around 140–160 m (Fig. 3). We

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FIG. 3. Leading covariance EOF of winter (DJF) 50-mb geopotential height for NNR1 data, and also for L38 and L60 NOVOLC experiments. Contour interval is 20 m, with gray areas indicating negative anomalies and dark gray areas indicating anomalies less than 2100 m.

can thus conclude with more than 95% confidence that over the last 150 yr (Stenchikov et al. 2006). The surface climate models are failing to capture a robust observed signal is characterized by anomalous warming of up to high-latitude response to volcanic aerosol. 2 K over northern Europe (north of 508N in Fig. 4) and anomalous cooling that reaches 21 K over southern b. Extratropical response to volcanic aerosol forcing Europe (south of 508N). Figure 4 shows that the L38 and The observed AO response to the three volcanic L60 models capture a statistically significant strength- eruptions is illustrated in Fig. 4. A significant strengthening of the polar vortex with a peak Z50 magnitude ening of the winter polar vortex is seen in the NNR1 around 70%–75% of that observed and centered just off analysis, with Z50 anomalies reaching 2240 m near the the pole. The observed surface signal is also broadly North Pole. This represents an anomalously positive reproduced by each model with statistically significant phase of the AO and is characteristic of the stratospheric anomalies over both northern and southern Europe. response to low-latitude volcanic eruptions observed The positive AO signal in the models is further seen in

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FIG. 4. The 50-mb (left) geopotential height (Z50) and (right) surface temperature anomalies for winter (DJF) averaged over six winters with volcanic forcing relative to climatology for NNR1, L38, and L60. Stippling indicates statistically signiﬁcant anomalies at the 90% conﬁdence level using a two-tailed t test (Spiegel 1988). For Z50, the contour interval is 20 m with gray areas indicating negative anomalies and dark gray areas indicating anomalies less than 2100 m. For temperature, contour interval is 0.5 K with gray areas indicating positive anomalies and dark gray areas indicating anomalies greater than 1 K.

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50- and 850-mb zonal wind, mean sea level pressure, and winter response to volcanic aerosol forcing alone. This is surface precipitation anomalies for both models (not despite the fact that the models demonstrate a realistic shown), highlighting the fact that the models adequately lower-stratospheric signal associated with aerosol heat- reproduce the observed signal. We also note that both ing at low latitudes (Figs. 4 and 6). model configurations give a similar response: differences d. Impact of initial and boundary condition forcing in the strength of the European AO signal between the models are not statistically significant. We there- We finally attempt to further understand the impact of fore conclude that increasing the vertical resolution in initial and boundary condition forcing on the extra- the stratosphere and extending the model lid into the tropical winter signal in the models by examining the mesosphere does not significantly affect the AO/NAO December–February signals separately. Figure 7 pres- response to volcanic aerosol forcing in our seasonal ents monthly mean Z50 anomalies in NNR1 and each hindcasts. model, composited over all postvolcanic years (relative to climatology), with error bars representing 90% con- c. Extratropical response without volcanic aerosol fidence intervals based on model ensemble spread using forcing a two-tailed t test. The anomaly for December in each The winter response of our initialized hindcasts to model is near 2200 m, similar to that observed. If this volcanic forcing is not consistent with the IPCC AR4 signal was being forced into the model through the climate models evaluated by Stenchikov et al. (2006), surface (SST) boundary we would expect it to persist which tended to produce only a weak positive AO re- over the entire winter. However, the strength of the sponse. A major difference here is that our model is run model anomaly is halved in January and the signal is in seasonal hindcast mode while those in recent studies completely gone in February, while the observed re- were run as long climate simulations. We thus consider sponse remains strong through the winter. (We note that the possibility that the AO signal captured by our models small differences between the observed and modeled may in fact originate from initial conditions and surface monthly anomalies may arise from the larger number of boundary forcing (SST) anomalies that contain the ob- years used in the observed climatology.) This result served signal, rather than from the implementation of strongly suggests the role of initial conditions in forcing volcanic aerosol itself. We test this by repeating the the model with the observed extratropical response to analysis of the previous section except for the NOVOLC volcanic eruptions at the beginning of the hindcast pe- experiments relative to climatology. These experiments riod. This is further supported by an atmosphere-only therefore only include the influence of initial and long climate simulation using the L60 model, which also boundary condition forcing on the extratropical winter shows a weak postvolcanic extratropical signal despite circulation, with the absence of volcanic aerosol. Sur- being forced with both volcanic aerosol and observed prisingly, the stratospheric polar and surface European SSTs (not shown). anomalies in Fig. 5 are similar to those shown in Fig. 4, The persistence of the AO signal in the first month of revealing that the strong intensification of the polar the hindcasts is consistent with the 30–35-day radiative vortex in the model hindcasts is in fact due to either time scale in the lower stratosphere of the northern persistence of the observed signal present in the initial polar region during winter (Kiehl and Solomon 1986; conditions or forcing from observed SSTs. The two fig- Newman and Rosenfield 1997). Moreover, previous ures only differ near the tropics where the presence of studies have shown that, while statistical forecasts of volcanic aerosol forcing in the lower stratosphere leads northern European winter surface anomalies decrease to an anomalous Z50 increase resulting from enhanced in skill with increasing lead time, there is an increase in radiative warming. skill at 30–35 days when the forecasts include strato- For completeness we present the difference between spheric information (e.g., Christiansen 2005). The re- Figs. 4 and 5; a composite of postvolcanic experiments sults of our study thus lend support to the premise that relative to the NOVOLC experiments to see the impact seasonal winter forecasts are substantially improved of volcanic aerosol forcing alone, without the influence with the inclusion of lower-stratospheric anomalies of initial/boundary condition forcing (since the initial (Charlton et al. 2003; Christiansen 2005). and boundary conditions are identical for postvolcanic and NOVOLC experiments). The weak stratospheric 4. Summary, discussion, and perspective polar and surface European signals seen in Fig. 6 con- firm that, consistent with the IPCC AR4 climate simu- The extratropical winter response to volcanic aerosol lations in Stenchikov et al. (2006), our model hindcasts forcing and its sensitivity to stratospheric resolution in fail to adequately capture the observed extratropical seasonal hindcasts are investigated in two versions of

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FIG. 5. The 50-mb (left) geopotential height (Z50) and (right) surface temperature anomalies for winter (DJF) averaged over six winters for NOVOLC experiments relative to climatology for L38 and L60. Stippling indicates statistically significant anomalies at the 90% confidence level using a two-tailed t test (Spiegel 1988). For Z50, the contour interval is 20 m, with gray areas indicating negative anomalies and dark gray areas indicating anomalies less than 2100 m. For temperature, contour interval is 0.5 K with gray areas indicating positive anomalies and dark gray areas indicating anomalies greater than 1 K. the Met Office Hadley Centre climate model with dif- and extending the model domain to near the mesopause. ferent vertical resolutions above the tropopause. Sea- The presence of this signal in the models, however, is sonal hindcasts initialized on 1 December produce a likely due to the persistence of the observed signal from strengthening of the winter polar vortex and anomalous the initial conditions, with the models showing a weak warming over northern Europe characteristic of the response to volcanic aerosol forcing alone that is consis- positive phase of the Arctic Oscillation (AO) when tent with previous studies on the impact of volcanic forced with volcanic aerosol following the 1963 Mount forcing in long climate simulations (e.g., Stenchikov et al. Agung, 1982 El Chichoń, and 1991 Mount Pinatubo 2006). The persistence of the observed signal in the first eruptions, as is observed. The AO signal in the L60 ex- month of the hindcasts is consistent with the 30–35-day tended model is of comparable strength to that in the L38 radiative time scale in the lower stratosphere and sup- standard model, suggesting that there is little impact from ports the premise that seasonal winter forecasts are sub- both increasing the vertical resolution in the stratosphere stantially improved with the inclusion of stratospheric

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FIG. 6. The 50-mb (left) geopotential height (Z50) and (right) surface temperature anomalies for winter (DJF) averaged over six winters with volcanic forcing relative to NOVOLC experiments for L38 and L60. Stippling indicates statistically significant anomalies at the 90% confidence level using a two-tailed t test (Spiegel 1988). For Z50, the contour interval is 20 m with gray areas indicating negative anomalies and dark gray areas indicating anomalies less than 2100 m. For temperature, contour interval is 0.5 K with gray areas indicating positive anomalies and dark gray areas indicating anomalies greater than 1 K. information. Hence, the models are able to simulate the in lower-stratospheric tropical heating, the extratropical response to volcanic eruptions if initialized in early win- winter response remained weak, suggesting a failure ter, but they do not demonstrate extratropical circulation to generate a strong feedback between the enhanced sensitivity to volcanic aerosol forcing. equator-to-pole temperature gradient and circulation It is presently not clear why climate models show a anomalies at high latitudes. Future work to untangle the weak extratropical response to volcanic aerosol. In the possible reasons behind these model deficiencies could early stages of this work we examined the possibility include (i) implementing a more realistic spatial structure that the polar stratosphere is sensitive to changes in the of volcanic aerosol that varies smoothly in latitude, lon- vertical distribution of volcanic aerosol in the tropical gitude, and time (particularly in the vicinity of the polar stratosphere (not shown). Postvolcanic sensitivity exper- night jet); and (ii) taking account of the volcanically in- iments (relative to NOVOLC experiments) showed that duced ozone depletion that was observed in the mid- while the models reasonably simulated associated changes latitudes following the El Chichoń and Mount Pinatubo

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waves, such that the model may be overrepresenting wave propagation into the polar stratosphere leading to a destabilization of the postvolcanic winter vortex. Poten- tial biases in the planetary wave response are supported by an out-of-phase relationship between the modeled and observed standing waves, as is suggested by the nonzonal component of the response seen in Fig. 4. The likely importance of planetary wave refraction is also supported by the absence of signiﬁcant winter anomalies in the observed annular mode of the Southern Hemisphere where planetary wave activity is weak and the strong polar vortex remains resistant to perturbations forced by volcanic aerosols (Robock et al. 2007). Investigation of this issue also requires further detailed analysis.

Acknowledgments. The authors thank Gareth Jones at the Met Office Hadley Centre, and Manoj Joshi and Keith Shine at the University of Reading for helpful discussions during the course of this work. We also thank the three anonymous reviewers for suggested improve- FIG. 7. Monthly mean Z50 anomalies for December–February averaged over six winters with volcanic forcing relative to clima- ments to the manuscript. ERA-40 data used in this study tology for NNR1, L38, and L60. Error bars represent 90% confi- were provided by ECMWF, and the NCEP–NCAR dence intervals based on model ensemble spread using a two-tailed reanalyses were provided by the NOAA/CIRES Climate t test (Spiegel 1988). Diagnostics Center. This work was supported by the Joint DECC and MoD Integrated Climate Programme (DECC) GA01101, (MoD) CBC/2B/0417_Annex C5. eruptions (e.g., Hofmann and Solomon 1989; Angell 1997; Solomon 1999), which may intensify the polar vortex through high-latitude radiative cooling (Stenchikov et al. REFERENCES 2002). Note however that the observed AO response to Alley, R. B., and Coauthors, 2007: Summary for policymakers. the nine largest volcanic eruptions from 1860 to 1999 [of Climate Change 2007: The Physical Science Basis, S. Solomon which 7 occurred prior to 1980 (Stenchikov et al. 2006)] et al., Eds., Cambridge University Press, 18 pp. is of similar strength to the observed response to the two Andronova, N. G., E. V. Rozanov, F. Yang, M. E. Schlesinger, post-1980 eruptions alone. The possible role of volca- and G. L. Stenchikov, 1999: Radiative forcing by volcanic nically induced ozone depletion observed after 1980 in aerosols from 1850 to 1994. J. Geophys. Res., 104, 16 807– 16 826. strengthening the AO response to volcanic forcing is Angell, J. K., 1993: Comparison of stratospheric warming following therefore unlikely to explain the full discrepancy be- Agung, El Chicho´ n and Pinatubo volcanic eruptions. Geo- tween models and observations. phys. Res. Lett., 20, 715–718. Despite the model deficiencies seen here, HadGEM1 ——, 1997: Estimated impact of Agung, El Chicho´ n and Pina- produces realistic AO variability (Fig. 3) and a realistic tubo volcanic eruptions on global and regional total ozone after adjustment for the QBO. Geophys. Res. Lett., 24, 647– AO response to ENSO, the QBO, and stratospheric sud- 650. den warming events (Marshall and Scaife 2009; Marshall ——, and J. Korshover, 1964: Quasi-biennial variations in tem- and Scaife 2009, manuscript submitted to J. Geophys. perature, total ozone, and tropopause height. J. Atmos. Sci., Res.). The model’s failure to respond to volcanic aerosol 21, 479–492. forcing, however, may suggest an issue regarding the ——, and ——, 1978: Estimate of global temperature variations in the 100–30 mb layer between 1958 and 1977. Mon. Wea. Rev., high-latitude climatology of the model; specifically, the 106, 1422–1432. propagation of planetary waves from the troposphere Baldwin, M. P., and Coauthors, 2001: The quasi-biennial oscilla- into the polar stratosphere, which can influence the tion. Rev. Geophys., 39, 179–229. high-latitude dynamical response to volcanic forcing Baran, A. J., and J. S. Foot, 1994: A new application of the oper- (Kodera 1994; Perlwitz and Harnik 2003). Characteris- ational sounder HIRS in determining a climatology of sul- phuric acid aerosol from the Pinatubo eruption. J. Geophys. tics of wave propagation will be influenced by subtle Res., 99, 25 673–25 679. differences between observed and modeled mean cli- Charlton, A. J., A. O’Neill, D. B. Stephenson, W. A. Lahoz, and mates or by a model bias in the generation of planetary M. P. Baldwin, 2003: Can knowledge of the state of the

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