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Model Projections of the Changes in Atmospheric Circulation and Surface Climate over North America, the North Atlantic, and Europe in the Twenty-First Century NGAR-CHEUNG LAU AND JEFFREY J. PLOSHAY NOAA/Geophysical Fluid Dynamics Laboratory, Princeton University, Princeton, New Jersey (Manuscript received 13 March 2013, in final form 21 June 2013) ABSTRACT The impacts of climate change on the North America–North Atlantic–Europe sector are studied using a coupled general circulation model: the Climate Model, version 3 (CM3) and a high-resolution atmosphereonly model, the High Resolution Atmospheric Model (HiRAM)—both developed at the Geophysical Fluid Dynamics Laboratory. The CM3 experiment is conducted under two climate change scenarios for the 1860– 2100 period. The sea surface temperature (SST) forcing prescribed in the ‘‘time slice’’ integrations with HiRAM is derived from observations for the 1979–2008 period and projection by CM3 for the 2086–95 period. The wintertime response in the late twenty-first century is characterized by an enhancement of the positive phase of the North Atlantic Oscillation in sea level pressure (SLP) and poleward and eastward displacements of the Atlantic jet stream and storm track. The forcing pattern due to eddy vorticity fluxes in the perturbed storm track matches well with the response pattern of the SLP field in the late twenty-first century. The model results suggest that the above circulation changes are linked to the gradient of the altered SST forcing in the North Atlantic. In summer, the projected enhancement of convection over the eastern tropical Pacific is accompanied by a wave train spanning the North America–North Atlantic–Europe sector. This quasi-stationary circulation pattern is associated with diminished storm track activity at 408–508N and an eddy forcing pattern similar to the summertime SLP response in the late twenty-first century.
1. Introduction A rich variety of meteorological features with multiple temporal and spatial scales prevails over the North Atlantic basin and the surrounding land areas. Particularly noteworthy phenomena in this region include the quasi-stationary jet stream extending from North America to the North Atlantic; synoptic-scale disturbances that preferentially travel along the zonally oriented ‘‘storm track’’; and a prominent mode of variability known as the North Atlantic Oscillation (NAO) (see Vallis and Gerber 2008; Hurrell and Deser 2009), which is associated with notable fluctuations in the location and intensity of both the Atlantic jet stream and storm track. Within the North Atlantic oceanic domain, the near-surface thermal conditions are intimately linked to the principal current systems such as the Gulf Stream
Corresponding author address: Jeffrey J. Ploshay, NOAA/ Geophysical Fluid Dynamics Laboratory, Princeton University, Forrestal Campus, 201 Forrestal Road, Princeton, NJ 08540. E-mail:
[email protected] DOI: 10.1175/JCLI-D-13-00151.1
and the North Atlantic and Labrador Currents, as well as deep-water formation near the northern terminus of the Atlantic meridional overturning circulation (AMOC) (see reviews by Kuhlbrodt et al. 2007; Lozier 2012; Srokosz et al. 2012). A considerable degree of coupling occurs among the myriad phenomena in the atmosphere– ocean system. For instance, the strength and location of the atmospheric jet stream and storm track are influenced by sea surface temperature (SST) anomalies owing to variability of the ocean currents or intensity of oceanic convection (e.g., Brayshaw et al. 2009). Conversely, changes in the atmospheric flow patterns could affect the energy and water fluxes across the air–sea interfaces and thereby modulate the characteristics of the underlying ocean sites (e.g., Woollings et al. 2012). Through various downstream influences, fluctuations in the circulation pattern over North America and the coupled air–sea system over the North Atlantic have strong implications on the response of surface atmospheric conditions over Europe to long-term climate changes. In particular, salient shifts in the intensity and position of the prevalent atmospheric surface westerly
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flow across North America and the North Atlantic could change the surface air temperature distribution over Europe through advective effects. Variations in the storm tracks originating from North America and North Atlantic could also have noticeable effects on the precipitation pattern over Europe. Numerous observational and modeling studies have been conducted to enhance our understanding of the impacts of the perturbed loadings of greenhouse gases and other trace constituents on the climate system in the North America–North Atlantic–European region. For the sake of brevity, we shall only mention a small subset of representative investigations that are directly relevant to the findings to be presented in this article. A majority of the model projections in various climate scenarios anticipate a gradual weakening of the AMOC through the twenty-first century (Meehl et al. 2007), which may partially be attributed to decreased nearsurface water density in the high-latitude North Atlantic region. The reduced oceanic convection and poleward heat transport in that region exerts a notable influence on the local SST pattern, especially during the winter season. Of particular interest is the large modelprojected cooling trend of the surface waters in the vicinity of the southern tip of Greenland (Dai et al. 2005; Stouffer et al. 2006a,b). Some of the model experiments also project intensification as well as northward or eastward displacements of the atmospheric storm tracks over the North Atlantic (Bengtsson et al. 2006; Ulbrich et al. 2008; Brayshaw et al. 2009; McDonald 2011; Catto et al. 2011; Woollings et al. 2012). Analysis by Stephenson et al. (2006) reveals that a large fraction of the models participating in phase 2 of the Coupled Model Intercomparison Project (CMIP2) yields a positive trend in the strength of the NAO in response to greenhouse gas forcing. The models contributing to that study are also in agreement with regards to future changes in the surface winter climate over Europe, with a warming trend over the entire continent and precipitation increases over northern Europe. The primary objective of the present study is to investigate in further detail the various interactions among the key atmospheric and oceanic phenomena in the North America–North Atlantic–Europe region within the context of climate change forced by different emission scenarios. This goal is achieved by diagnosing a suite of coordinated model experiments, which was conducted at the Geophysical Fluid Dynamics Laboratory (GFDL) in support of the Fifth Coupled Model Intercomparison Project. The model runs considered in this investigation range from an ensemble of multicentury simulations based on a fully coupled atmosphere–ocean general circulation model (GCM) with moderate spatial resolution,
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to integrations with an atmosphere-only GCM with much higher resolution for selected periods in the late twentieth and twenty-first centuries. Our attention will be focused on the local and downstream effects of changes in the North Atlantic SST and North American circulation patterns on the storm track and dynamical interactions between transient disturbances traveling along the storm track and the ambient circulation pattern. The seasonal dependence of some of the above relationships will be explored by analyzing the model output for both winter and summer seasons. The design of various model experiments examined in this study is described in section 2. The changes in atmospheric circulation and surface climate over North America–North Atlantic–Europe in different seasons, as projected by a high-resolution atmospheric GCM forced by a range of climate scenarios for the late twenty-first century, are documented in section 3. The relationships between projected changes in the wintertime SST pattern in the North Atlantic, and the summertime circulation pattern over North America, on the strength and position of the storm track are investigated in section 4. The detailed evolution of the projected circulation changes are explored in section 5 using output from a coupled GCM experiment covering the entire twenty-first century. The principal findings are summarized and discussed in section 6.
2. Model tools and experimental design The primary model tools for this study were developed at GFDL and consist of the following GCMs: d
d
a fully coupled climate model, the GFDL Climate Model, version 3 (hereafter CM3). Details of the atmospheric component of this model (AM3) are described by Donner et al. (2011). The model atmosphere has a horizontal resolution of approximately 200 km. Vertical variations are represented in 48 layers. New physical parameterizations have been incorporated in AM3 to treat aerosol–cloud interactions and chemistry–climate interactions. The oceanic component of CM3 is based on the Modular Ocean Model (MOM4p1) code of Griffies (2009) and the parameterization schemes described by Gnanadesikan et al. (2006). The horizontal resolution of the model ocean is 18 in latitude and longitude poleward of 308. Resolution in the meridional direction is increased in the tropics, reaching 1/ 38 at the equator. Vertical variations are represented in 50 layers. The treatment of land hydrology and sea ice processes in CM3 is documented in the appendix of Donner et al. (2011). a high-resolution atmosphere-only model, the High Resolution Atmospheric Model (hereafter HiRAM).
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The formulation of this model is described by Zhao et al. (2009). The finite-volume dynamical core uses a cubed-square grid topology (Putman and Lin 2007). Results will be presented for two versions of this model, with horizontal resolutions of approximately 50 km (referred to as the C180 version) and 25 km (the C360 version). There are 32 atmospheric layers in the vertical direction. The cloud fraction is computed using a diagnostic scheme assuming a subgrid-scale distribution of total water. Convective processes are parameterized on the basis of the shallow convection scheme of Bretherton et al. (2004). The primary database for the present study is generated in various experiments with the above models. These model runs include the following: d
d
integration of CM3 through the sequence of climate scenarios in the 1860–2100 period. In the historical era (1860–2005), the emission data (for aerosols and ozone) and prescribed concentrations (for greenhouse gases) are based on observational estimates. From 2006 onward, these forcings are assumed to follow the representative concentration pathways (RCP) issued by the Intergovernmental Panel for Climate Change (see van Vuuren et al. 2011). Separate integrations are performed for the RCP4.5 and RCP8.5 scenarios, with total radiative forcing reaching the level of 4.5 and 8.5 W m22 by 2100, respectively. Five parallel runs are completed for the historical period of 1860–2005. For the RCP4.5 (RCP8.5) scenario, three (one) parallel runs are conducted for the 2006–2100 period. ‘‘time slice’’ runs with the HiRAM model for simulating the current climate in the period of 1979–2008 (referred to as the NOW experiment) and for projecting the future climate in the period of 2086–95 ( the PROJ experiment). In the NOW experiment, the lower boundary conditions at oceanic grid points are prescribed using intermonthly variations of observed SST and sea ice based on the first Hadley Centre Sea Ice and Sea Surface Temperature dataset (HADISST1) for the 30-yr duration of those experiments. The model atmosphere is also subjected to temporally varying radiative forcings, as estimated from concentrations of aerosols and different trace gases prescribed for the 1979–2008 epoch. In the PROJ experiment, separate runs are conducted using radiative forcings based on concentrations of various trace constituents corresponding to the RCP4.5 and RCP8.5 scenarios (see van Vuuren et al. 2011). The SST condition and sea ice concentration are computed by taking the difference between the mean SST and sea ice concentration fields in the 2086–95 and 1979–2008 periods, as generated in the CM3 experiment for a given climate
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scenario, and then adding this projected SST and sea ice change to the HADISST1 data for individual years in the 1999–2008 period. Using the HiRAM model with C180 resolution, three parallel runs are completed for the NOW experiment, and for the PROJ experiments subjected separately to RCP4.5 and RCP8.5 forcings. In addition, two parallel runs of the NOW experiment and three runs of the PROJ experiment for the RCP8.5 scenario are performed using the C360 version of the HiRAM model.
3. Seasonal dependence of projected changes in atmospheric circulation and surface climate The atmospheric response of the HiRAM model to prescribed radiative and SST forcings corresponding to various climate scenarios is illustrated in Fig. 1 for the December–February (DJF) season and in Fig. 2 for the June–August (JJA) season. The projected changes in the sea level pressure (SLP, left panels) and 250-mb zonal wind (right panels) fields are computed by subtracting the ensemble mean of the output of the NOW experiment from that of the PROJ experiment for various scenarios. Throughout the presentation in this article, the ensemble means for a given experiment are obtained by averaging over data for all available years in all parallel runs conducted for that experiment. In Figs. 1 and 2, the differences between the PROJ and NOW experiments are shown using shading only at those grid points where such differences exceed the 95% significance level based on a two-tailed Student’s t test. In the right panels, the location of the North Atlantic jet stream in the present climate, as obtained from the ensemble mean of the 250-mb zonal wind data generated in the NOW experiment, is indicated using contours. Comparison of the contour patterns in the right panels of Figs. 1 and 2 with their observational counterparts [e.g., see Hoskins et al. (1989) or the 40-yr European Centre for Medium-Range Weather Forecasts Re-Analysis (ERA-40) atlas accessible at http://www.ecmwf.int/ research/era/ERA-40/ERA-40_Atlas/docs/index.html] indicates that present-day climatological strength and position of the North Atlantic jet stream are well simulated by the HiRAM model. For both RCP4.5 and RCP8.5 scenarios and in both the C180 and C360 versions, the HiRAM model yields a wintertime SLP pattern (Figs. 1a–c) characterized by rising SLP along a zonal belt extending from the central North Atlantic to the Mediterranean region and falling SLP over northern Canada and the Greenland/ Norwegian Seas. The large-scale features of this pattern bear some correspondence to those associated with the positive phase of the observed NAO (e.g., Hurrell and
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FIG. 1. Distributions of the difference between the PROJ (2086–95) and NOW (1979–2008) experiments with the HiRAM, for the climatological (left) sea level pressure and (right) 250-mb zonal wind fields (shading) in the winter (DJF) season. Results are based on integrations with the C180 version subjected to the (a),(d) RCP4.5 and (b),(e) RCP8.5 scenarios and (c),(f) with the C360 version subjected to the RCP8.5 scenario. The contour patterns superposed on (right) depict the present-day climatology of the 250-mb zonal wind field, as obtained from the NOW experiment with the corresponding model version. Red or blue shading is shown only over grid points where the difference between the PROJ and NOW experiments surpasses the 95% significance level, as determined by a two-tailed Student’s t test. High-terrain areas with orographic heights exceeding 1500 m are indicated by gray shading.
Deser 2009). However, the primary negative center projected over northeastern Canada is shifted to the west of the observed NAO center near Iceland, whereas the positive center in the model patterns (especially that projected by C360, see Fig. 1c) is displaced to the east of the observed location. The patterns in Figs. 1a–c indicate strengthening of the positive phase of the NAO toward the end of the twenty-first century. The model response to the RCP8.5 scenario (Fig. 1b) is noticeably stronger than
that to the RCP4.5 scenario (Fig. 1a). The projections based on the C180 and C360 versions are qualitatively similar to each other, with the amplitudes of signals in the C180 response (Fig. 1b) being higher than those in the C360 response (Fig. 1c) at some locations. The most prominent feature in the corresponding patterns for the 250-mb zonal wind (shading in the right panels of Fig. 1) is the increased wind speeds in an elongated zone extending northeastward from the
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FIG. 2. As in Fig. 1 but for the summer (JJA) season.
western United States to northern Europe. Particularly noteworthy is the intensified westerly flow over the North Atlantic at 508–558N. The latter changes are located to the northeast of the axis of the time-mean jet stream in the present climate (see contours in Figs. 1d–f). These results are hence indicative of a tendency for the North Atlantic jet to shift poleward and to penetrate farther downstream into Europe in the course of the twenty-first century. The projected wind changes are larger in the RCP8.5 scenario (Fig. 1e) than in the RCP4.5 scenario (Fig. 1d). The C360 version (Fig. 1f) produces stronger wind signals over North America than the C180 version (Fig. 1e).
The spatial scale and typical amplitudes of the SLP responses in the northern summer (left panels of Fig. 2) are smaller than those in the winter season. The positive summertime SLP center over the North Atlantic is displaced poleward of its winter counterpart by 58–108 latitude. The overall pattern in the Atlantic sector, Figs. 2a–c, is similar to that of the NAO during the summer season (e.g., see Hurrell and Deser 2009). The right panels of Fig. 2 indicate that the pattern of zonal wind changes at 250 mb in the summer months exhibits a banded structure, with increasing wind speeds at 558–658N and decreasing speeds within a
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FIG. 3. Variations with latitude (ordinate) and calendar month (abscissa) of the difference between the PROJ and NOW experiments, for the partial zonal average of climatological sea level pressure over the longitudes spanning eastward from 408W to 408E. Results are based on integrations with the C180 version subjected to the (a) RCP4.5 and (b) RCP8.5 scenarios and (c) with the C360 version subjected to the RCP8.5 scenario. Positive (negative) differences that are significant at the 95% level are indicated by bold solid (dashed) black contours.
broader belt between 358 and 508N (with most notable speed reduction over North America). These bands of opposing polarity straddle the mean axis of the summertime jet stream over North America and North Atlantic simulated in the NOW experiment (see contours in Figs. 2d–f). This spatial relationship signifies poleward displacement of the summer jet stream during the twentyfirst century, as well as reduction of the wind speed near the climatological jet stream axis.
A more detailed depiction of the seasonal dependence of the dipolar pattern of SLP change in the North Atlantic–Europe region is presented in Fig. 3, which shows the partial zonal averages of the SLP field from 408W to 408E, where the north–south SLP seesaw is particularly strong (see left panels of Figs. 1, 2). Results are obtained by subtracting the ensemble means of the NOW experiment from the corresponding PROJ experiment subjected to various scenarios.
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The patterns in Fig. 3 illustrate that the strongest SLP changes are projected in the winter season with statistically significant increases (decreases) at 358–458N (658–808N). For the C180 version subjected to RCP4.5 scenario (Fig. 3a), this dipolar feature does not meet the 95% significance criterion in the summer. The same summertime feature is more discernible in both the C180 and C360 projections under the RCP8.5 scenario (Figs. 3b,c), with the midlatitude positive center being shifted poleward of its wintertime position by 108–158 latitude. The projected changes in the European surface climate associated with the atmospheric circulation responses noted in Figs. 1–3 are summarized in Fig. 4. Differences in the surface air temperature (SAT) and precipitation fields between the PROJ and NOW experiments based on the C180 model are presented here. The patterns based on experiments with the C360 version (not shown) are almost the same as their counterparts obtained from the C180 version. In view of the strong similarity between the projections by the C180 and C360 versions of the temperature, precipitation, and circulation fields (see also Figs. 1–3), we shall henceforth focus on products of the C180 version only and not present any more results from the C360 version. During the winter season, the patterns for precipitation changes (Figs. 4a,b) indicate increased dryness over the Iberian Peninsula and parts of the Mediterranean region, where rising SLP is projected (see Figs. 1a–c). Conversely, wetter conditions prevail over most European sites situated poleward of 508N. These locations are under the influence of intensified surface westerlies, as inferred from the SLP patterns in Figs. 1a–c using geostrophy. The charts for SAT (Figs. 4c,d) show statistically significant warming throughout the entire domain considered here. The strongest temperature rises are projected over northern Russia, whereas more modest warming is anticipated over the Mediterranean region and Western Europe. Such regional contrasts in the magnitude of the SAT trend may partially be attributed to temperature advection by the near-surface circulation changes accompanying the SLP patterns in Figs. 1a–c. Specifically, strengthening of the southerly flow over northern Russia, which lies to the east of the deepened cyclone center over the Greenland–Norwegian Seas, may accentuate the warming in the Russian sector. On the other hand, the warming near the Mediterranean and Western Europe is alleviated by enhanced northerly flows over that region, located east of the intensified anticyclone over North Atlantic. In the summer season, results from the PROJ experiment under the RCP8.5 scenario indicate significant drying and enhanced warming at most European sites
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south of 558N, as well as increased precipitation and comparatively less warming at locations north of 608N (Figs. 4f,h). The prominent poleward shift of the drying zone from winter to summer is consistent with the corresponding seasonal displacement of the positive SLP center, as noted in Figs. 1–3. Increased subsidence in the lower troposphere (not shown) prevails over the summertime drying zone between 408 and 558N (Fig. 4f).
4. SST conditions, quasi-stationary circulation patterns, and storm track dynamics associated with climate changes We proceed to investigate the mechanisms related to the long-term trends in atmospheric circulation and surface climate, as documented in the previous section. We shall focus our attention on the patterns of changes in SST forcing and the background circulation, as well as their relationships with perturbed eddy forcing induced by the storm tracks.
a. Winter season The distributions of the SST change in the North Atlantic from the 1979–2008 period to the 2086–95 period, as projected by the coupled model CM3, are shown in Fig. 5 for the RCP4.5 and RCP8.5 scenarios. As explained in section 2, these SST perturbations are incorporated as the lower boundary forcing for the PROJ experiment based on the atmospheric HiRAM model. Figure 5 indicates rising temperatures at most grid points in the North Atlantic, with the notable exception of the region centered near 558N, 328W and another smaller site east of Iceland. This SST pattern is in general agreement with the corresponding result generated by a previous version of the coupled GCM at GFDL, as presented by Stouffer et al. (2006a). These authors and other investigators pointed out that the SST cooling trend at the high-latitude sites in the North Atlantic is the manifestation of reduced poleward oceanic heat transport and downwelling of seawater associated with a weakened Atlantic meridional overturning circulation (AMOC). Between 258 and 408W the meridional SST gradient in the vicinity of 508N is strengthened since this site is straddled by SST cooling to the north and warming to the south. The close relationship between this dipolar pattern in SST change and AMOC weakening has also been noted by Zhang (2008) and Zhang et al. (2011). However, the evidence presented by Winton et al. (2013) indicates that the retardation of the AMOC as projected by CM3 is noticeably stronger than the corresponding decline in other CMIP models developed outside of GFDL. The higher sensitivity of the AMOC response in CM3 to climate change could lead to
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FIG. 4. Distributions of the difference between the PROJ and NOW experiments, for the climatological (top) precipitation and (bottom) surface air temperature fields. Results for the (left) winter (DJF) and (right) summer (JJA) seasons are based on integrations with the C180 version subjected to the (a),(c),(e),(g) RCP4.5 and (b),(d),(f),(h) RCP8.5 scenarios. Shading (see color scale) is shown only over grid points where the difference between the PROJ and NOW experiments surpasses the 95% significance level.
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FIG. 5. Distributions of the difference between the PROJ and NOW experiments, for the climatological (left) prescribed sea surface temperature and (right) thickness between 850 and 700 mb (contours) and Eady growth rate (shading) in the winter (DJF) season. Results are based on integrations with the C180 version subjected to the (a),(c) RCP4.5 and (b),(d) RCP8.5 scenarios.
exaggeration of the SST cooling trend off the southeastern coast of Greenland (see Figs. 5a,b). Various observational and modeling studies have demonstrated that SST signals in the North Atlantic are linked to atmospheric patterns throughout the entire troposphere—in the context of climatological data (Minobe et al. 2008), month-to-month variability (Nakamura and Yamane 2009), and individual storms (Booth et al. 2012). In the present study, we focus on the association of SST change with the thermal structure of the lower troposphere, as illustrated by the thickness between the 850-mb and 700-mb surfaces (see contours in Figs. 5c,d). This thickness field provides for a measure of the averaged air temperature between the two pressure surfaces. The patterns for this field are dominated by a zonally elongated minimum with an axis at ;558N.
The center of this minimum is located just to the east of the principal site of SST cooling in North Atlantic (cf. the left and right panels of Fig. 5). The increased meridional atmospheric temperature gradient at 508N over the central and eastern portions of North Atlantic is evident from the contour patterns in Figs. 5c,d. A convenient measure of the baroclinic instability of the mean flow environment, as formulated by Eady (1949) and applied by Hoskins and Valdes (1990) and James (1994), is the maximum growth rate s 5 0.31fDU/NH. Here f is the Coriolis parameter, DU the vertical wind shear (linked to meridional temperature gradient by the thermal wind relationship), N the square root of Brunt– V€ ais€ al€ a frequency, and H the atmospheric scale height. In the present case, DU is computed as the difference between the zonal wind at 700 and 850 mb. The values of s
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FIG. 6. Distributions of the difference between the PROJ and NOW experiments, for the climatological (left) rms (shading) of the bandpass (2.5–6 days) filtered height and (right) eddy-induced height tendency, all at 250 mb during the winter season (DJF). The contour patterns superposed on (left) depict the present-day climatology of the rms of bandpass filtered 250-mb height, as obtained from the NOW experiment. Results are based on integrations with the C180 version subjected to the (a),(c) RCP4.5 and (b),(d) RCP8.5 scenarios.
are obtained separately from the ensemble means of the wind data from the NOW and PROJ experiments. The differences in this parameter for the two experiments are depicted using shading in Figs. 5c,d. It is seen that the Eady growth rate is projected to increase along a zonal belt centered at ;508N. This belt is located south of the elongated minimum in the thickness field (see contours in Figs. 5c,d) and is hence characterized by enhanced baroclinicity. Conversely, decreased growth rates are projected at 608–708N where reduced baroclinicity prevails. We proceed to examine the actual changes in storm track behavior associated with the trends in baroclinicity of the mean flow environment. Following the methodology of Blackmon (1976), the amplitude and location of synoptic-scale eddy activity are inferred from the rootmean-square (rms) of geopotential height fluctuations at 250 mb with periods between 2.5 and 6 days, as retained by application of a Lanczos bandpass filter (Duchon 1979). The differences in this rms amplitude between the PROJ and NOW experiments are shown, using shading, in Fig. 6 for the RCP4.5 and RCP8.5 scenarios. The distribution of the rms field based on the NOW experiment is mapped in Figs. 6a,b using contours. The location of
the present-day climatological storm tracks, as inferred from the axis of the maximum in these contour patterns, is in good agreement with observations [e.g., see Lau et al. (1981) or the ERA-40 atlas]. Inspection of the shaded patterns in Figs. 6a,b reveals that storm track activity is projected to increase most noticeably along an elongated zone centered at 508–608N and stretching from the central North Atlantic to northern Europe. Over the central North Atlantic, this feature lies in the vicinity of increased Eady growth rates (see shading in Figs. 5c,d) and intensified zonal flow at 250 mb (shading in right panels of Fig. 1). The spatial relationship between the shaded and contour patterns in Figs. 6a,b is indicative of a trend for northward and downstream extension of the storm track from its present-day location. In the European sector, the enhanced storm track activity corresponds to an increase from the present-day local climatological rms values (contours in Figs. 6a,b) of approximately 25% (40%) in the RCP4.5 (RCP8.5) scenario. The projected precipitation increases in the European sector north of 508N (Figs. 4a,b) may partially be attributed to the more active synoptic-scale disturbances over that region.
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The above relationship between changes in SST gradient and storm track position/intensity is in accord with the findings presented by Brayshaw et al. (2009) and Woollings et al. (2012), based on diagnoses of the output from the Hadley Centre Climate Model, version 3 (HadCM3) and the suite of models contributing to the CMIP3, respectively. Both of these studies reported poleward and eastward displacements of the storm track in the North Atlantic when the strength of AMOC is reduced. The transient disturbances embedded in the midlatitude storm tracks are active agents of vorticity transport (e.g., see Holopainen 1978; Lau and Wallace 1979). Dynamical forcing of the background flow may be expressed in terms of the convergence of such eddy fluxes. We now examine the nature of eddy forcing by computing the geopotential height tendencies induced by the eddy vorticity fluxes at 250 mb: ›Z f 5 =22 [2$(V0 z0 )] . ›t eddy g Here g is gravitational acceleration, V the horizontal velocity vector, and z relative vorticity. The overbar represents time average, and the prime represents deviation of the bandpass filtered fluctuations from this average. The vorticity flux can be evaluated using the eddy statistics u0 u0 , y 0 y 0 , and y 0 u0 (see Lau and Holopainen 1984). The climatological average of (›Z/›t)eddy , is computed using the output from the PROJ and NOW experiments separately, and the difference in this quantity between the two experiments is displayed in the right panels of Fig. 6. These panels show that negative eddyinduced height tendencies are projected over much of the Arctic zone poleward of 558N, whereas positive height tendencies are evident along a zonal belt lying to the south of that latitude. Comparison between the shaded patterns in the left and right panels of Fig. 6 indicates that these two regions of opposing height tendencies straddle the axis of enhanced storm track activity in the North Atlantic–Europe sector. This spatial organization of eddy forcing about the perturbed storm track axis is consistent with the dynamical arguments of Hoskins et al. (1983). In view of the barotropic nature of the forcing due to eddy vorticity fluxes (e.g., see Lau and Nath 1991), the pattern of (›Z/›t)eddy at the sea level is qualitatively similar to that at 250 mb (Figs. 6c,d). The spatial correspondence between the principal features in Figs. 6c,d and those in the SLP charts (Figs. 1a–c) suggests that uppertropospheric storm track dynamics could contribute to the near-surface circulation pattern response to long-term climate change.
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b. Summer season The patterns for the difference between the JJA climatologies of the PROJ and NOW experiments are shown in Fig. 7 for the 250-mb height (contours) and precipitation (shading) fields and for the RCP 4.5 and RCP 8.5 scenarios. The axis of the climatological intertropical convergence zone (ITCZ), as deduced from the belt of maximum precipitation appearing over the tropical Pacific and Atlantic Oceans in the NOW experiment, is depicted on these charts using broad pink bands. In both scenarios, the 250-mb height response is characterized by a wave train spanning across the North America–North Atlantic–Europe domain, with alternating cyclonic (blue contours) and anticyclonic (red contours) centers. This wave train emanates from the eastern tropical Pacific, where a prominent wetting trend is projected along a zonal belt at about 58N (green shading). The wet belt is shifted slightly equatorward of the climatological ITCZ axis (broad pink band). The linkage of the extratropical wave train with precipitation changes over the tropical Atlantic appears to be weaker than that with rainfall perturbations over the eastern tropical Pacific. The seasonal pattern of precipitation change over the eastern tropical Pacific and its spatial relationship with the local ITCZ are in accord with the findings presented by Huang et al. (2013) based on zonal averages produced by 18 individual CMIP5 models. The latter authors noted that two mechanisms are at play in determining the location of maximum precipitation changes. The first is the ‘‘wet get wetter’’ effect (Held and Soden 2006), which leads to precipitation increases in climatologically wet regions, such as the Pacific ITCZ. The second is the ‘‘warmer get wetter’’ effect (Xie et al. 2010), which yields more precipitation over regions where the SST rise exceeds the mean surface warming in the tropics. In the Pacific sector, enhanced warming is projected by CM3 (not shown) to occur at the equator owing to weakening of the Walker circulation and surface easterlies (Liu et al. 2005; Vecchi and Soden 2007). The placement of the maximum precipitation change between the ITCZ and the equator may hence be interpreted as the net result of these two effects. In the tropical Atlantic sector, CM3 projects strongest SST warming at about 108N (not shown) where the local climatological ITCZ is situated. Hence both the warmerget-wetter and wet-get-wetter effects lead to increased precipitation near 108N, with almost no latitudinal shift between the projected precipitation trend (shading in Fig. 7) and the climatological ITCZ (broad pink band in Fig. 7) over the Atlantic. Trend patterns analogous to Fig. 7 have also been constructed using output from the PROJ and NOW
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FIG. 7. Distributions of the difference between the PROJ and NOW experiments, for the climatological 250-mb height (with red and blue contours indicating cyclonic and anticyclonic flows, respectively) and precipitation (shading), all for the summer season (JJA). Results are based on integrations with the C180 version subjected to the (a) RCP4.5 and (b) RCP8.5 scenarios. Position of the ITCZ, as deduced from the climatological precipitation pattern based on the NOW experiment, is indicated by broad pink bands.
experiments for the DJF season. The wintertime results (not shown) indicate noticeably weaker changes in precipitation and SST in the eastern tropical Pacific as compared with the corresponding JJA responses. Hence, we shall focus on the atmospheric changes associated with the wave train pattern during the summer season. In the North American sector, the wave train response in 250-mb height during the JJA season (contours in Fig. 7) consists of a cyclonic belt at 258–408N and an anticyclonic belt at 408–558N. A strong increase in near-surface air temperature (not shown) is projected underneath the anticyclone. This warm-core high center is accompanied by reduced eastward wind speeds in the
358–508N zone and increased wind speeds farther north (see also Figs. 2d–f). The changes in storm track characteristics and eddy forcing in conjunction with the altered background state in summer are described in Fig. 8 using rms of bandpass filtered geopotential height (Figs. 8a,b, shading) and (›Z/›t)eddy (Figs. 8c,d), all at the 250-mb level. The climatological rms statistics based on output from the NOW experiment are superposed in Figs. 8a,b using contours. These contour patterns compare well with their counterparts based on observational data [e.g., see Lau et al. (1981) or the ERA-40 atlas], thus indicating realistic simulation by HiRAM of the mean location of the summertime storm track in present-day
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FIG. 8. As in Fig. 6 but for the summer season (JJA).
climate. A prominent belt of reduced synoptic-scale activity is seen to extend from western North America to Europe within the 358–508N zone. Parallel to this feature is a band of increased rms values centered near 658N. The placement of this pair of features relative to the mean axis of the storm track in the present-day climate (see contours in Figs. 8a,b) is again suggestive of a poleward shift of the center of transient activity during the summer season. This northward displacement of the storm track is linked to the upper-tropospheric anticyclonic center in the 408–608N zone over North America (Figs. 7c,d), which tends to divert synoptic disturbances toward higher latitudes. The spatial quadrature relationship between the patterns for changes in rms amplitude and (›Z/›t)eddy , as noted previously for winter (Fig. 6), is also discernible in the summer season (Fig. 8). In particular, the axis of reduced eddy activity at about 458N is straddled by positive height tendencies to its north and negative tendencies to its south. Conversely, negative (positive) height tendencies are projected in the region lying to the north (south) of the belt of enhanced rms amplitudes at about 658N. The patterns of eddy forcing in Figs. 8c, d bear some spatial correspondence with those of the summertime changes in 250-mb height (Figs. 7c,d) and
SLP (Figs. 2a–c), thus suggesting dynamical feedbacks between the high-frequency and slowly varying components of the atmospheric circulation.
5. Corresponding results based on experiments with CM3 The changes projected by CM3 for the RCP 4.5 scenario are shown in Fig. 9 for SLP (Figs. 9a,d), rms of bandpass filtered 250-mb height (Figs. 9b,e), and (›Z/›t)eddy (Figs. 9c,f). The corresponding patterns for the RCP8.5 scenario (not shown) are similar to those displayed in Fig. 9, except that the amplitudes of the projected signals for RCP8.5 are relatively higher. In the winter season, CM3 projects positive (negative) changes in SLP and (›Z/›t)eddy near 458N (658N) over the North Atlantic–European sector (Figs. 9a,c) and increased storm track activity along a belt centered at 508–608N (Fig. 9b). In the summer season, the principal centers of SLP changes (Fig. 9d) are shifted poleward of their wintertime positions; storm track disturbances are weakened along 358–458N and strengthened along 508–658N (Fig. 9e); and (›Z/›t)eddy is characterized by negative tendencies near 358N and positive tendencies near 458N (Fig. 9f). Most of the above features exhibit
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FIG. 9. Distributions of the difference between the climatologies of the CM3 experiment, computed separately for the 2086–95 and 1979–2008 periods, for (a),(d) sea level pressure; (b),(e) rms bandpass-filtered 250-mb height; and (c),(f) eddy-induced height tendency at 250 mb. Results are based on output for the (left) winter (DJF) and (right) summer(JJA) seasons and for the RCP4.5 scenario.
some correspondence with their counterparts deduced from the HiRAM model, as presented in sections 3 and 4. Contrary to the time-slice approach of the NOW and PROJ experiments using the HiRAM model, the CM3 runs cover the entire span from 1860 to 2100. The
output from the latter runs hence offers an opportunity to examine the detailed temporal evolution of various features throughout the twentieth and twenty-first centuries. In Fig. 10 the distributions with respect to time and latitude of partial zonal average of SLP and rms
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FIG. 10. Variations with latitude (ordinate) and year (abscissa) of the partial zonal average of sea level pressure between (a) 58E and 258E in winter, (b) 458W and 08 in summer, and of rms bandpass-filtered 250-mb height between (c) 808W and 258E in winter and (d) 908W and 08 in summer. Results are based on ensemble means for individual years in the GFDL CM3 experiment subjected to the RCP4.5 scenario. The longitudinal ranges used to compute the partial zonal averages of various quantities are designed to capture the phenomena of interest (see Figs. 9a,d for the SLP field and Figs. 9b,e for the rms of 250-mb height field). All values represent departures from the average of the corresponding quantity over the 1971–2000 period. An 11-yr running mean is applied to the yearly time series.
of bandpass filtered 250-mb height for selected longitudinal sectors in both seasons are shown. The SLP patterns for both seasons (Figs. 10a,b) are characterized by the emergence of a dipolar signal during
the early decades of the twenty-first century, with positive changes south of 558N (608N) in winter (summer) and negative changes farther north. The amplitude of such changes exhibit decadal-scale fluctuations and
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attain peak values in the middle and late parts of the twenty-first century. The patterns of rms 250-mb height are dominated by an enhanced level of synoptic-scale activity at 508–558N in winter (Fig. 10c) and reduced activity at 408–458N in summer (Fig. 10d). The magnitude of these changes increases monotonically with time throughout the twenty-first century so that maximum changes in storm track intensity are projected toward the end of the period being considered here. Most of the panels in Fig. 10 exhibit relatively weak signals in the twentieth century and prominent trends in the twenty-first century. The sharp contrasts between the evolution in the two epochs may be partially linked to the insufficient warming of the global-mean surface air temperature in the CM3 simulation for the twentieth century [as compared to observations, see Golaz et al. (2013)] and the excessive warming in the CM3 projection for the twenty-first century [as compared to results from a previous version of the climate model—the GFDL CM2.1; see Levy et al. (2013)]. The reliability of the trends appearing in the CM3 experiment needs to be tested further by diagnosing output from other models.
6. Summary and discussion The response in the North America–North Atlantic– European sector to projected changes in radiative and SST forcing in the late twenty-first century is investigated by subjecting a high-resolution atmospheric GCM to these conditions. In the winter season, the response pattern in the SLP field is characterized by a zonal belt of positive changes along 408–458N and negative change over the Arctic zone (Figs. 1a–c). Moreover, the intensity of the upper-level zonal flow is enhanced poleward and downstream of the present-day climatological axis of the North American jet stream (Figs. 1d–f). In the summer season, the belt of positive SLP change is shifted noticeably northward of its wintertime position (Figs. 2a–c), and the zonal flow is weakened (strengthened) on the equatorward (poleward) flank of the presentday jet stream (Figs. 2d–f). These circulation changes exert noticeable influences on the precipitation and surface air temperature over many parts of Europe in both seasons (Fig. 4). The relationship between the wintertime response and the SST changes in the North Atlantic is explored. The SST pattern is characterized by a cooling trend off the southeastern coast of Greenland (Figs. 5a,b). The enhanced meridional SST gradient along the zonal belt lying to the south of this region is accompanied by strengthened baroclinicity in the lower troposphere (Figs. 5c,d) and more vigorous storm track activity over
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the North Atlantic (Figs. 6a,b). The dynamical feedback of the altered storm track on the background flow is illustrated using eddy-induced geopotential tendencies (Figs. 6c,d). This pattern of eddy forcing exhibits a spatial correspondence with that of the projected SLP change (Figs. 1a–c). In the summer season, the response in the upper troposphere is dominated by a wave train pattern emanating from the tropical Pacific, with prominent anticyclonic centers along the belt between 458 and 558N (Fig. 7). This wave train is apparently linked to the precipitation trend over the eastern Pacific at about 58N, due to a combination of the ‘‘wet get wetter’’ and ‘‘warmer get wetter’’ effects. The altered flow environment in the Western Hemisphere is conducive to weakened synoptic-scale activity to the south of the present-day storm track and increased activity to the north (Figs. 8a,b). The pattern of eddy-induced height tendencies associated with these storm track changes (Figs. 8c,d) again bears a correspondence with the upper-tropospheric geopotential height and SLP response patterns in summer (Figs. 7a,b and 2a,c). The results presented in this article are based on model tools developed at GFDL. The robustness of the inferences drawn from these results needs to be assessed by analogous detailed analysis of the simulations performed by other modeling groups contributing to CMIP5. For instance, considering that the AMOC response to climate change in CM3 is stronger than that in most CMIP models (Winton et al. 2013), it would be of interest to determine whether the findings on the relationships between wintertime changes in SST, storm track activity, and eddy forcing (see section 4a) are supported by experimentation with other models. Other issues that are worthy of further investigations include (i) the relative roles of various oceanic features and processes (e.g., AMOC, current systems, sea ice melting, surface heat fluxes) in generating SST changes in different parts of the North Atlantic, (ii) the atmospheric responses in the North America–Europe sector to near-local boundary conditions (e.g., extratropical SST, land surface characteristics) and to more remote forcings (e.g., tropical SST and precipitation), and (iii) the nature of dynamical interactions between the storm track and background flow field in different climate change scenarios. Acknowledgments. We thank our colleagues at GFDL for providing the output from various experiments with the C180, C360, and CM3 models. We are also indebted to Larry Horowitz, Ronald Stouffer, Rong Zhang, and the official reviewers for their perceptive comments on an earlier version of this manuscript.
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