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  • View in gallery

    Annual cycle of area-weighted average NA SWE (solid lines) and snow extent (SNE; dashed lines) for high- and low-snow forcing.

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    Seasonal SWE (black contour lines) and surface temperature (color-filled contours) response to EY snow forcing during (a) fall, (b) winter, (c) spring, and (d) summer. Positive SWE contours (solid lines) drawn at 0.5, 5, 15, and 25 cm. Negative SWE contours (dashed lines) drawn at −0.5, −1, −2, and −3 cm. Only statistically significant (>95%) surface temperatures are plotted.

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    Three-month running mean 500-hPa geopotential height variance ()−1/2 response to EY snow forcing. Seasons begin with fall (SON), and progress, ending with ASO. Contours drawn at ±1, 2, 3, and 4 m; solid (dashed) contours denote a positive (negative) response. 90% (95%) statistical significance indicated with light (dark) gray shading.

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    Spring season transient activity response to EY snow forcing. (a) 500-hPa EKE; contours drawn at ±1, 2, 3, and 4 m2 s−1. (b) 750-hPa meridional temperature flux ; contours drawn at ±0.25°, 0.75°, 1.0°, and 1.5°C m s−1. (c) 750-hPa vertical temperature flux ; contours drawn at 0.01°K Pa s−1 intervals. (d) 750-hPa EGR; contours drawn at ±0.02, 0.04, 0.08, and 0.12 day−1. Solid (dashed) contours denote a positive (negative) response. 90% (95%) statistical significance indicated with light (dark) gray shading.

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    Spring season spatial gradient response to EY snow forcing at 750 hPa. (a) Vertical shear −∂u/∂p; contours drawn at 0.001 m s−1 intervals. (b) Meridional temperature gradient |∂T |/∂y; contours at ±4°, 6°, 8°, 10°, 12°, 16°, 20°, and 24°C 1000 km−1. Solid (dashed) contours denote a positive (negative) response. 90% (95%) statistical significance indicated with light (dark) gray shading.

  • View in gallery

    Spring season latitude–pressure profile response to EY snow forcing averaged over NA longitudes (140°–60°W). (a) Zonal wind u; contours drawn at ±0.5, 1, 1.5, and 2 m s−1. (b) Air temperature T; contours drawn at ±0.5°, 1°, 1.5°, and 2°C. Solid (dashed) contours denote a positive (negative) response. 90% (95%) statistical significance indicated with light (dark) gray shading.

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    Column-averaged diabatic heating response to EY forcing during (a) winter, and (b) spring. Contours drawn at ±0.5 K day−1 intervals. Solid (dashed) contours denote a positive (negative) response. 90% (95%) statistical significance indicated with light (dark) gray shading.

  • View in gallery

    As in Fig. 2, but for FS snow.

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    As in Fig. 3, but for FS snow.

  • View in gallery

    Seasonal 750-hPa meridional temperature gradient |∂T |/∂y response to (a) EY and (b) FS snow forcing. (left) Fall, (middle) winter, and (right) spring seasons. Contours drawn at ±4°, 6°, 8°, 10°, and 12°C 1000 km−1. Solid (dashed) contours denote a positive (negative) response. 90% (95%) statistical significance indicated with light (dark) gray shading.

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    Winter 850-hPa stationary wave streamfunction response to FS forcing. Contours drawn at ±0.5, 0.75, 1.0, and 1.25 × 106 m2 s−1. Solid (dashed) contours denote a positive (negative) response. 90% (95%) statistical significance indicated with light (dark) gray shading.

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    The 500-hPa stationary wave streamfunction ensemble mean (EY low snow) during (a) winter and (c) spring; contours drawn at ±2 × 106 m2 s−1 intervals. 500-hPa stationary wave streamfunction response to EY forcing during (b) winter and (d) spring; contours drawn at ±0.5 × 106 m2 s−1 intervals. Solid (dashed) contours denote a positive (negative) response. 90% (95%) statistical significance indicated with light (dark) gray shading.

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Modeled Climate State and Dynamic Responses to Anomalous North American Snow Cover

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  • 1 Columbia University, New York, New York
  • | 2 Lamont Doherty Earth Observatory, Columbia University, Palisades, New York
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Abstract

The radiative and thermal properties of widespread snow cover anomalies have the potential to modulate local and remote climate over monthly to seasonal time scales. In this study, physical and dynamical links between anomalous North American snow conditions and Northern Hemisphere climate are examined. A pair of 40-member ensemble AGCM experiments is run, with prescribed high- and low-snow forcings over North America during the course of an entire year (EY). The difference between the two ensemble averages reflects the climatic response to sustained EY snow forcing. Local surface responses over the snow forcing occur in all seasons, and a significant remote surface temperature response occurs over Eurasia during spring. A hemispheric-scale transient eddy response to EY forcing also occurs, which propagates downstream from the forcing region to Eurasia, and then reaches a maximum in extent and amplitude in spring. The evolution of this transient eddy response is indicative of considerable downstream development and is consistent with known storm-track dynamics. This transient response is shown to be a result of persistent steepened temperature gradients created by the anomalous snow conditions, which contribute to enhanced baroclinicity over the storm-track entrance regions. A second pair of experiments is run, with the prescribed high- and low-snow forcings over North America restricted to the fall season (FS). The dynamical response to FS forcing is muted compared to the EY scenario, suggesting that the seasonal timing and persistence of the snow forcing are essential for the remote teleconnection.

Corresponding author address: Stefan Sobolowski, 500 West 120th Street, Dept. of Earth and Environmental Engineering, Columbia University, New York City, NY 10027. Email: sps2109@columbia.edu

Abstract

The radiative and thermal properties of widespread snow cover anomalies have the potential to modulate local and remote climate over monthly to seasonal time scales. In this study, physical and dynamical links between anomalous North American snow conditions and Northern Hemisphere climate are examined. A pair of 40-member ensemble AGCM experiments is run, with prescribed high- and low-snow forcings over North America during the course of an entire year (EY). The difference between the two ensemble averages reflects the climatic response to sustained EY snow forcing. Local surface responses over the snow forcing occur in all seasons, and a significant remote surface temperature response occurs over Eurasia during spring. A hemispheric-scale transient eddy response to EY forcing also occurs, which propagates downstream from the forcing region to Eurasia, and then reaches a maximum in extent and amplitude in spring. The evolution of this transient eddy response is indicative of considerable downstream development and is consistent with known storm-track dynamics. This transient response is shown to be a result of persistent steepened temperature gradients created by the anomalous snow conditions, which contribute to enhanced baroclinicity over the storm-track entrance regions. A second pair of experiments is run, with the prescribed high- and low-snow forcings over North America restricted to the fall season (FS). The dynamical response to FS forcing is muted compared to the EY scenario, suggesting that the seasonal timing and persistence of the snow forcing are essential for the remote teleconnection.

Corresponding author address: Stefan Sobolowski, 500 West 120th Street, Dept. of Earth and Environmental Engineering, Columbia University, New York City, NY 10027. Email: sps2109@columbia.edu

1. Introduction

Anomalous continental-scale snow cover can influence both local and downstream climate because of its radiative and thermal properties, which act to modify the overlying atmosphere (e.g., Barnett et al. 1989; Cohen and Entekhabi 2001; Cohen and Rind 1991; Leathers and Robinson 1993). These influences may occur from regional to hemispheric spatial scales and immediate to seasonal time scales. Snow cover is a seasonally varying land surface state that covers much of the Northern Hemisphere from mid- to high latitudes. Because of this spatiotemporal variation and the striking geographical differences between the North American and Eurasian landmasses, the influence of snow on climate and the physical pathways through which this influence is expressed is still an area of ongoing research. In particular, questions regarding large-scale dynamic responses and their corresponding mechanisms abound.

Evidence for physically based snow climate teleconnections is more abundant for Eurasia than North America (NA). This is likely due to the greater Eurasian landmass, the existence of well-known centers of stationary wave activity over Siberia and the Tibetan Plateau (Plumb 1985), and thus greater potential for large-scale snow anomalies to influence the overlying atmosphere. Gong et al. (2003a) reported a hemispheric response to anomalous Siberian snow cover that resembles the Arctic Oscillation (AO). This response is constrained by the unique topography of the region, such that the snow-enhanced stationary wave activity propagates upward and poleward (Gong et al. 2004). Others noted a possible link between Eurasian snow cover and the North Atlantic Oscillation (NAO; Qian and Saunders 2003; Saunders et al. 2003) or the Indian monsoon (Bamzai and Shukla 1999), and evaluated its potential for climate predictability (Cohen and Entekhabi 2001). Recent studies questioned the ability of atmospheric general circulation models (AGCMs) to capture observed Eurasian snow–northern annular mode teleconnections, and noted that the modeled response may be highly dependent on model-specific treatment of troposphere–stratosphere coupling (Hardiman et al. 2008).

Relatively little attention has been paid to the dynamical response of the atmosphere to NA snow conditions. Gong et al. (2003b) described a weak positive AO-like response to a fall snow forcing. Sobolowski et al. (2007, hereafter SGT07) identified a more robust response involving downstream stationary wave fields and winter surface temperatures over central Eurasia. Klingaman et al. (2008) suggested a link between anomalous Great Plains snow cover and wintertime western Eurasian surface air temperatures facilitated by a snow-induced shift in the NAO toward positive phase. Conversely, G. Henderson et al. (2008, personal communication) reported a positive Pacific–North American pattern response, a weakened NAO-like response, and a central-eastern Eurasian winter temperature response to anomalous NA snow cover. These studies exhibit some qualitative similarities, but also clear differences and inconsistencies, which indicate that the dynamics of a snow-forced atmospheric response to NA snow cover is still not well understood.

The immediate local–regional influence of snow cover on NA surface temperatures is well documented, and it is associated with radiative, sensible, and latent heat fluxes (Ellis and Leathers 1998; Groisman et al. 1994; Leathers and Robinson 1993; Namias 1985). Another localized effect of anomalous snow cover may be seen in the overlying atmosphere as diabatic cooling enhances sinking, anticyclonic motion, and, thus, changes in air mass (e.g., Ellis and Leathers 1998; Leathers et al. 2002; Serreze et al. 1998).

The general climate state over NA during the winter and transition seasons is largely determined by the position and magnitude of the stationary and quasi-stationary waves and their associated transient disturbances. The wintertime stationary wave field itself is determined by tropical and extratropical diabatic heating, orography, transient eddies, and nonlinear interactions (Chen and Trenberth 1988a,b; Chen 2005; Held et al. 2002; Ting et al. 2001). Held et al. (2002) noted that there are still outstanding questions regarding the forcing of the wintertime stationary waves, which may be complicated by the two-way nature of the interactions between forcings and the waves themselves. In particular, the precise nature of the interactions between planetary waves and transient eddies that collectively make up the Northern Hemisphere storm tracks is still unclear [see Chang et al. (2002), for a review of storm-track dynamics]. The situation is complicated by the fact that the storm tracks themselves likely owe their existence, at least in part, to orography and heating, as well as variations of the planetary wave field (Broccoli and Manabe 1992; Hoskins and Valdes 1990; Lee and Mak 1996).

Research conducted to date indicates that the atmospheric response to anomalous NA snow is largely confined to the troposphere and may best be described as modulating the existing time mean circulation patterns. This is in contrast to the response to Eurasian snow cover, which suggests a troposphere–stratosphere pathway. Different physical mechanisms to those suggested for the response to Eurasian snow cover are likely at work over NA. However, specific mechanisms by which snow conditions initiate and maintain such a response are not readily apparent. Further, given the dependence of midlatitude transient eddies on thermal forcing, a broad thermal forcing such as that represented by anomalous snow conditions may elicit a transient response. We seek to address these gaps in our knowledge by investigating the AGCM-simulated climate response to NA snow forcing with a focus on the high-frequency transient response and associated North Atlantic storm tracks. The aim of this study is to better define the mechanisms/pathways by which NA snow anomalies may influence local and remote climate, and the effects that snow anomaly timing and duration have on this response. The modeling experiments are detailed in section 2. Simulation results are presented in sections 3 and 4. A discussion of the physical mechanisms that generate the simulated transient wave response, and its possible interaction with stationary wave activity, follows in section 5. Conclusions are posited in section 6, which consolidates our findings.

2. Methods

a. ECHAM5 AGCM

The ECHAM5 AGCM is the latest-generation global climate model developed by the Max Planck Institute for Meteorology and used in the Intergovernmental Panel on Climate Change Fourth Assessment Report (Roeckner et al. 2003). The ECHAM5 model has been shown to adequately simulate the observed radiative fluxes and hydrological cycle (Hagemann et al. 2006; Roesch and Roeckner 2006; Wild and Roeckner 2006). Roesch and Roeckner (2006) assessed ECHAM5’s treatment of land surface albedo, snow cover fraction, and snow depth. Even though the surface albedo parameterization is still oversimplified, simulation of the seasonal cycle of snow depth and the timing of seasonal snowmelt is improved over the previous ECHAM4 model, which overestimated Eurasian snow mass in the spring because of delayed snowmelt. The annual cycle, interannual variability, and late-twentieth-century trends of snow-covered area are also simulated reasonably well in ECHAM5. We run ECHAM5 with climatological sea surface temperature and sea ice boundary conditions, at T42 spatial resolution (roughly 2.8°) to capture broad hemispheric continental-scale phenomena with reasonable computational efficiency.

b. Snow-forcing specifications

We aim to create realistic but idealized continental-scale snow forcings, such that snow extent/depth are always, and everywhere, larger (smaller) than climatology for a positive (negative) forcing. We make use of a recently constructed gridded NA snow depth dataset that extends from 1900 to 2000 (Dyer and Mote 2006). This dataset is constructed from daily observations at 7000 stations, which are then interpolated to a 1° × 1° grid using a procedure for irregularly spaced data. An independent quality review found that the data quality decreases dramatically prior to 1956, so we only make use of the 45-yr period from 1956 to 2000 (Ge and Gong 2008). The dataset extends from 25° to 75°N for the entire zonal extent of NA.

Monthly mean snow depths over the period of record are computed for grid cells with a snow cover frequency ≥50%, otherwise snow depth is set to 0 (snow frequency = days with snow per days in month). These monthly grid cell values are used to determine the maximum and minimum NA spatial snow extent over the period of record for each calendar month, and their corresponding snow lines. This serves to spatially constrain the model snow forcings to the observed envelope of snow-covered regions. The associated monthly snow depth fields are determined by selecting the maximum and minimum snow depth over the period of record at each grid cell above the snow line. To keep the low-snow forcing small but not snow free, a lower limit of 2.5-cm snow depth is applied throughout. What results are 24 gridded snow depth fields representing realistic high- and low-snow-forcing conditions (with respect to both extent and depth) for each of the 12 calendar months.

To prepare these fields for input into ECHAM5, snow depth is converted to snow-water equivalent (SWE), using a conservative value of 100 kg m−3 for the average snow density, and the 1° × 1° snow depth grid is scaled to the T42 model resolution. To ensure that these snow forcings consistently portray positive (negative) anomalies, for each month the snow lines for the two forcings are compared to make sure that they do not cross, and the converted maximum (minimum) gridcell SWE values are checked to make sure they exceed the models climatological maximum (minimum) gridcell SWE. Figure 1 shows the resulting spatially averaged SWE and extent of the prescribed monthly snow forcings, and indicates that the maximum SWE differences occur in late winter and early spring (February–April), while the maximum differences in extent occur during the transition seasons (May and October).

ECHAM5 runs on a 360-day year that consists of 12 months each, with six 5-day pentads. The monthly mean snow forcings at each grid cell are applied to the third pentad of each month and used to linearly interpolate snow forcings for the remaining pentads. This results in SWE forcing over NA that varies from week to week. This degree of temporal resolution is sufficient because we seek to observe influences on the monthly to seasonal time scale.

c. Experiment design

The experimental approach is to compare two 40-member ensemble AGCM simulations, which prescribe either the high or low pentad SWE fields described above. The integration period for each ensemble member is from September through August, with initial conditions obtained from a 40-yr control simulation with climatological sea surface temperatures and sea ice. At the beginning of each time step the prescribed SWE values are specified over NA grid cells. The response to an anomalous positive snow forcing is diagnosed by subtracting the low-snow ensemble average from the high-snow ensemble average. A variety of mean state and eddy variance statistics are calculated to assess the climate response to NA snow forcing. Where appropriate, Student’s t tests are used to assess statistical significance.

Two experiments are conducted. The first experiment prescribes the high- and low-snow forcings over the entire year’s integration period, and is designated EY. The second experiment prescribes the high- and low-snow forcings only during the fall [September–November (SON)] season, after which the models climatological SWE (obtained from the control simulation) is prescribed instead; this experiment is designated FS. The two experiments serve to distinguish the climatic response for multiseason versus fall-only NA snow forcings.

3. Climate response to EY

We divide the response to EY into three components. First, surface and near surface parameters are examined, both local to the snow forcing and in remote downstream regions. Next, we consider transient eddy activity in the atmosphere as a possible pathway for linking local and remote surface-state responses. Finally, local dynamical mechanisms that respond directly to snow forcing and can initiate and maintain an atmospheric pathway are investigated. Results are presented primarily as 3-month boreal seasonal mean responses to a positive snow forcing.

a. Local and remote state response

The SWE response for EY is shown in Fig. 2 as solid contour lines, which over NA provides a spatial representation of the magnitude and extent of the prescribed snow forcing. Over northern Eurasia a modest negative SWE response is apparent mainly during spring. The surface temperature response is also shown in Fig. 2, as filled contours. An immediate local response to the NA snow forcing is a pronounced diabatic cooling at the surface, resulting from the radiative and thermal effects of the increased snow cover and depth. This cooling is present throughout the snow season, and also during the summer, because of increases in soil moisture that result from melting the anomalous late winter snowpack (not shown). A significant remote surface warming, greater than or equal to 2.0°C, is also observed over northern Eurasia during the spring. A concurrent and collocated positive (negative) albedo response is present over NA (Eurasia), which is consistent with the SWE and temperature responses (not shown). The local surface responses are anticipated given the anomalous snow forcing; however, the robust downstream springtime temperature response over Eurasia is unexpected.

b. Storm-track/transient eddy response

While the negative SWE and albedo responses over Eurasia coincide with and are likely associated with the positive temperature response, it is not clear whether they initiate or arise from the temperature response. One possibility is that atmospheric transient eddy activity acts as a conduit by which a near-surface response over NA travels downstream, across the North Atlantic, and into Eurasia. Therefore, a variety of high-pass-filtered transient eddy statistics that collectively describe the storm-track zones over the Northern Hemisphere are computed, to assess their synoptic-scale response to the snow forcing. These zones include the classic storm-track regions over the North Atlantic and North Pacific, which may be better understood as existing embedded in bands of potentially high baroclinicity that circumnavigate the globe in the midlatitudes (e.g., Chang et al. 2002; Hoskins and Hodges 2002). To capture activity on a time scale of 2–8 days, a high-pass filter is applied to the daily departures from the monthly means of the relevant parameters before computing the transient eddy statistics (Trenberth 1991).

The midtroposphere (500 hPa) geopotential height variance response is shown in Fig. 3 as 3-month running means, which exhibit a clear and sustained intensification of transient eddy activity. It begins over and just downstream of the forcing region in the early fall, and propagates through the snow season to circumnavigate the globe, reaching its maximum extent in the springtime. The response is strongest over the NA forcing region and the primary North Atlantic storm-track region, but its considerable downstream propagation suggests that a signal originating in NA affects transient activity as far away as eastern Siberia and the North Pacific. This intensification of transient activity over regions outside of the traditional storm tracks and the propagation of the signal around the midlatitudes indicate significant downstream development (Chang 1993; Orlanski and Chang 1993; Simmons and Hoskins 1979). Additionally, as the snow line shifts southward from fall to winter the intensification of the transient eddy activity over NA is accompanied by a southward shift, as indicated by the negative (positive) dipole anomalies to the north (south).

Additional measures of transient eddy activity, which are physically consistent with known storm-track dynamics, support the intensification and propagation of storm-track activity seen in Fig. 3. These measures evolve similarly to the geopotential height variance from fall to spring, so only the maximum (spring) responses are shown here. The midtroposphere (500 hPa) eddy kinetic energy response (EKE) is written as
i1520-0442-23-3-785-e1
and is an indication of the transfer of available eddy potential to eddy kinetic energy as a result of baroclinic processes. Figure 4a shows a robust EKE response that follows a similar pattern to that of the geopotential height variance, that is, a pronounced southward shift and intensification over the forcing region and North Atlantic, and intensification across northern Eurasia extending to the North Pacific. The EKE response suggests that large amounts of energy are being transferred in the midtroposphere; the energetics of baroclinic instability theory require poleward and upward movement of heat for this to happen. Indeed, Figs. 4b,c show that both poleward and upward 750-hPa temperature fluxes exhibit increases over the same regions as the EKE and geopotential height variance responses. A modest 250-hPa momentum flux response indicates an intensified poleward flux around the storm-track exit regions that also is dynamically consistent with the other measures of transient eddy activity (not shown).
The measures of transient eddy activity presented here indicate an intensification of activity over both traditional and nontraditional storm-track regions. Given the enhanced poleward and upward movements of heat and the accompanying energy exchanges, these regions should also exhibit enhanced baroclinic growth. We compute the baroclinic growth as measured by the maximum Eady growth rate (EGR)
i1520-0442-23-3-785-e2
at 750 hPa, where N is the buoyancy frequency (Lindzen and Farrell 1980; Hoskins and Valdes 1990). The springtime EGR response shown in Fig. 4d exhibits enhanced baroclinicity over the United States, the North Atlantic storm-track entrance–exit regions, northern Eurasia, the North Pacific storm-track region, and the northwest coast of NA. As with the other measures of transient activity, the EGR response evolves and propagates through the snow season, reaching its maximum extent and amplitude in spring. The hemispheric pattern of the EGR response is consistent with the energetics suggested previously, that is, baroclinic theory and known storm-track dynamics. Additionally, the robust–coherent spatial and temporal structures of the transient eddy statistics suggests that growing eddies are sustained and that energy is recycled far downstream.

c. Snow-forced baroclinic enhancement

The nature of our snow forcing and the local surface temperature response suggests a modulation of the meridional temperature gradient |∂T |/∂y over NA from fall through spring. By the demands of thermal wind balance the dominant components of the EGR specification are linked to |∂T |/∂y and also vertical wind shear −∂u/∂p (Piexoto and Oort 1992; James 1995). Intensified |∂T |/∂y is also required to produce the modeled temperature fluxes. Thus, for the anomalous snow cover to initiate and sustain the transient eddy response, both −∂u/∂p and |∂T |/∂y must exhibit sustained intensification over the regions of maximum baroclinic growth and energy conversion (i.e., NA and the North Atlantic storm-track entrance regions).

Figure 5 shows that the time mean 750 hPa −∂u/∂p and |∂T |/∂y over NA both intensify and exhibit a dipole response with positive (negative) anomalies to the south (north). The temperature gradient dipole is also consistent with a mid- to high-latitude snow-forced cooling applied to an existing equator-to-pole temperature decrease, and is steepest to the south where the shifted storm tracks emerge. As with the previous results the temporal evolution is the same and the response in both fields is most pronounced in the springtime, over regions that also exhibit enhanced baroclinic growth. The meridional temperature gradient steepens along the south, east, and west coasts of NA and flattens in the northern interior. This represents the southward shift of the snow line in the high-snow case relative to the low-snow case.

Figure 6a shows a latitude–pressure plot of the springtime zonal wind response averaged over NA longitudes (140°–60°W). The zonal wind speed response increases with altitude in the lower troposphere over midlatitudes, consistent with the increase in 750-hPa −∂u/∂p shown in Fig. 5a. Figure 6a also indicates an intensification of the midlatitude jet, located in the upper troposphere between 30° and 50°N. This jet response is consistent with a positive transient eddy zonal mean flow feedback (e.g., Lau 1988; Lau and Holopainen 1984). Negative zonal wind responses both poleward and equatorward of the jet suggest that it may become more latitudinally confined and shifted poleward (because of the larger magnitude of the equatorward response) as it becomes stronger. Figure 6b shows a similar latitude–pressure plot of the springtime temperature response over the NA sector, which is confined to the lower troposphere. Increasing (decreasing) meridional temperature gradients are apparent near the surface around 30°N (50°N), consistent with Fig. 5b. These time mean responses follow the same seasonal evolution as the transient responses, reaching their maximum in springtime.

An important feature of the traditional midlatitude storm tracks is a region of enhanced diabatic heating in the troposphere along the western edge of the ocean basins that is coincident with growing synoptic storm systems and transient eddy activity (Hoskins and Valdes 1990). Figure 7 shows the column-averaged diabatic heating response over NA and adjacent oceans, for winter and spring. Diabatic cooling is apparent over much of NA during winter, which stems directly from the local snow forcing of surface air temperatures. Over the eastern United States, this surface cooling is masked by condensational heating above associated with enhanced baroclinicity. The resulting column-averaged diabatic heating response along the North Atlantic storm-track entrance region is evident in both winter and spring.

4. Climate response to FS

The climatic responses for the EY scenario reach their maximum magnitude and extent in the spring, but are apparent for most of the snow season. Because of the short time scale of transient eddies, this response pattern is likely due to the coinciding extreme magnitude of the EY depth forcing during late winter–early spring, and the persistence of the snow forcing over multiple seasons. The FS scenario aims to answer the following two questions: 1) is the seasonally constrained forcing in the FS scenario sufficient to produce a climate response, and 2) once the forcing is removed does the response persist or fall off? The impacts of magnitude and timing are important for understanding snow–climate interactions, especially in light of observed changes in seasonal snow cover variability and projected changes resulting from climate change (e.g., Brown 2000; Groisman et al. 1994; Raisanen 2008).

The SWE and surface temperature response for FS is shown in Fig. 8 as contour lines and filled contours, respectively. The fall SWE response over NA is identical to that of the EY scenario because the snow forcing is identical, while subsequent seasons show no response because the snow cover is held at climatology. The expected negative surface temperature response is seen over the forcing region during the fall, with only modest persistence of this local response into winter. FS indicates a positive remote springtime temperature response over the same general Eurasia region as for EY; however, the magnitude and coherence are reduced considerably. Examination of monthly plots (not shown) indicates that the FS spring Eurasian temperature response is almost entirely due to conditions in March, but the EY spring response is robust throughout the season. Thus, it is unclear if this FS spring temperature response is physically meaningful.

The geopotential height variance, shown in Fig. 9, exhibits a North Atlantic storm-track response from the fall to midwinter, but dissipates soon after. The FS response is qualitatively similar to the EY response shown in Fig. 3 up to early winter, and exhibits signs of downstream energy recycling. However, the FS response is clearly muted relative to the EY response, and dissipates quickly once the forcing is removed. As a result, the FS response does not propagate substantially downstream from winter to spring, and lacks the coherent hemispheric structure seen in Fig. 3 for EY.

Additional measures of transient eddy activity confirm a pattern of modest persistence of the FS signal into the winter months, which then falls off dramatically from winter into spring (not shown). The EKE response is quite robust in the fall season, but is mostly limited to the forcing region and the North Atlantic storm track, and all but disappears by spring. Likewise, the poleward and upward temperature fluxes exhibit a similar pattern but become even less coherent than EKE and the height variance as the snow season progresses. Overall, the transient activity metrics do not indicate a consistent and coherent propagation of the signal through winter spring.

The mechanisms for sustained baroclinic enhancement observed in the EY response are largely absent for FS. Vertical shear exhibits modest enhancement across the North Atlantic storm-track region, but the pattern is difficult to distinguish from climatological noise (not shown). The meridional temperature gradient response follows a similar pattern of reduced magnitude, as do the vertical profiles of (u) and (T).To illustrate the reduced winter–spring baroclinicity response in the FS scenario, Fig. 10 compares the seasonal evolution of the NA temperature gradient response for EY versus FS (note that the upper-right panel is identical to Fig. 5b). In both experiments, the identical fall snow forcing appears to be strong enough to generate an immediate local and downstream baroclinic response. For EY the snow forcing and the baroclinic response are maintained in the following seasons, and grow dramatically by spring. In contrast, for FS the snow forcing is removed beginning in December, and the baroclinic response diminishes rapidly. The results of the FS scenario suggest that a NA snow forcing that persists throughout winter and spring is essential for maintaining and amplifying a hemispheric-scale transient activity and remote climate-state response.

5. Discussion

a. Atmospheric teleconnection pathway involving transient activity

The AGCM simulation results presented above depict a snow-forced atmospheric teleconnection pathway that stretches from NA to eastern Siberia and the North Pacific via transient eddy activity. The immediate effect of the prescribed snow forcing is a pronounced diabatic cooling over NA. As a consequence, the meridional temperature gradient steepens in the southern portion of the forcing region while flattening to the north. There is also significant steepening of the temperature gradients along the coasts of NA (east, west, and south). These snow-forced temperature gradient responses are sufficient to generate increased poleward and upward temperature fluxes and hence enhanced baroclinic growth. This response occurs in the vicinity of the North Atlantic storm-track entrance region, and so translates into an intensified storm track. This storm track does not depict the movement of individual eddies, but rather is comprised of numerous eddies that grow, dissipate, and recycle their energy toward their downstream neighbors (Chang et al. 2002, and references therein). This downstream development, when combined with sustained baroclinic enhancement over the storm-track region, enables the transient response to a NA snow forcing to circumnavigate the globe.

An intensified and extended North Atlantic storm track can lead to the remote surface temperature response over northern Eurasia shown in Fig. 2c. The increased low-level meridional temperature flux (Fig. 4b) and negative sea level pressure responses (not shown) over this region suggests flow of warmer air into Eurasia. Thus, the snow-forced changes in meridional temperature gradients lead to enhanced baroclinicity and transient activity over the North Atlantic storm-track region, which extends into Eurasia and results in a flow pattern that leads to warmer late winter/early spring temperatures. This interpretation is supported by observation-based research that shows a link between North Atlantic storm-track intensification/extension and warmer Siberian temperatures (Rogers and Mosley-Thompson 1995; Rogers 1997).

b. EY versus FS forcing

The robust hemispheric transient response and its associated mechanisms are produced by the EY experiment, but largely disappear when the snow forcing is limited to the fall season in the FS experiment. This result suggests that the sustained EY snow forcing that persists through the entire snow season is necessary to yield the atmospheric teleconnection pathway and remote transient-state response in Eurasia during winter–spring. The responsiveness of the winter season is reasonable because the snow forcing migrates southward and is collocated with the North Atlantic storm-track entrance region. The storm tracks are most active during the winter season, so that the snow forcing works to amplify an already active background state. Further, winter–spring snow forcing can be considerable if snow depth is considered along with snow extent, as has been done here. Ge and Gong (2008) suggest that snow depth on its own may be of sufficient magnitude and scope to modulate regional hemispheric climate in a manner unrelated to extent. Figure 1 shows that the prescribed snow depth forcing actually peaks in early spring, which may help explain why the modeled transient response is strongest during the spring season. When the snow forcing is restricted to the fall season and high latitudes, both the forcing magnitude and the prevailing transient eddy activity are insufficient to sustain a hemispheric-scale atmospheric response.

Given the absence of the transient eddy teleconnection pathway for the FS experiment, it is unclear whether the fall snow forcing produces the apparent early spring Eurasian temperature response indicated in Fig. 8c. Note that the Siberian responses for FS are not entirely collocated with the Siberian response region for EY shown in Fig. 2c. One possible explanation is that another teleconnection pathway exists. The FS snow forcing occurs primarily at high latitudes (Figs. 2a and 8a), which may initiate an Arctic pathway that arrives in Eurasia in early spring. Another possibility is that the FS Eurasian temperature anomalies are simply climate noise, because the isolated responses in Fig. 8c lack the spatial and temporal coherence of the EY response (Fig. 2c). Without a discernable teleconnection mechanism, the remote temperature response for FS cannot be considered physically meaningful.

c. Comparison with SGT07

The FS experiment is similar to the basic experiment performed in SGT07, in that both model the climatic response to a fall snow forcing over NA. However, SGT07 utilized an older version of the ECHAM AGCM, had a different snow-forcing specification, and restricted the integration period to September–February. The snow forcing for SGT07 was taken directly from the 2 yr with the highest (lowest) fall snow observations over NA, unlike FS and EY, which take the highest (lowest) snow values over all years. Further, the snow forcing persisted into the winter months in SGT07 while it is cut off after November in FS. The SGT07 experiment yielded a modest North Atlantic storm-track response comparable to the FS fall–winter response shown in Fig. 3, which prompted storm-track anomalies to be cited as a possible causal mechanism in SGT07. This level of reproducibility between SGT07 and FS plus the physical pathway discerned for EY support the validity of the results presented here.

However, the main SGT07 result involving the winter stationary wave field and European climate-state responses to NA snow forcing is not reproduced by FS. As an example, Fig. 11 shows that the 850-hPa winter stationary wave response for FS is scattered, and does not portray the structured strengthening of the prevailing stationary wave field reported in SGT07. Unlike the modest transient response to a fall NA snow forcing, the stationary wave response in SGT07 is not reproduced by the FS experiment presented here. Note that mechanisms for the modeled response in SGT07 were postulated but not demonstrated. In contrast, a clear physical mechanism involving baroclinic enhancement and storm-track intensification is demonstrated for the EY experiment presented here.

d. Transient–stationary wave interactions

Even though FS does not reproduce the main SGT07 results, stationary waves may nonetheless be involved in the atmospheric teleconnection presented here. A stationary eddy response in the EY scenario is apparent but is not as robust as that of the transient eddies and their associated processes. Figure 12 shows the winter and spring 500-hPa stationary wave responses, alongside their prevailing ensemble mean fields for the low-snow case. In winter the stationary wave response is ambiguous, that is, the circulation pattern appears to strengthen over western Europe but weaken over the North Pacific (examination of the ensemble mean fields rules out significant eastward–westward shifting of the stationary waves). In spring the stationary wave response resembles the hemispheric pattern of the transient metrics. The weakening of the stationary waves over some regions (e.g., Pacific–NA) is consistent with known stationary–transient wave interactions (e.g., Lau and Holopainen 1984; Valdes and Hoskins 1989). Overall, it is not clear whether the stationary wave response is an indirect effect of snow-forced transient activity, a direct response to the NA snow forcing, or some combination of the two. Planetary stationary waves are known to be forced by diabatic heating, transient activity, and orography. The hemispheric storm tracks and transient eddy activity in general, owe their existence in part to orographic, stationary eddy, and diabatic heating forcings. Therefore, a full understanding of the atmospheric response to anomalous snow forcing, which is shown here to influence diabatic heating–cooling, and thereby transient activity, must also account for interactions between the transient and stationary eddy response. A full decomposition of the stationary wave response into its component forcings is underway to address this gap in our understanding.

6. Conclusions

This study presents the results from a suite of AGCM experiments designed to explore the climate impacts of anomalous NA snow forcing. Two experiments are performed: EY is subject to extreme but realistic snow forcings for the entire year, while FS is only subject to these snow forcings during the fall season. Each experiment is comprised of a high- and low-snow case, which are differenced. The major findings are summarized as follows:

  • (i) The EY snow forcing results in significant surface cooling over the forcing region in all seasons. Downstream, significant surface warming is seen over northern Eurasia in spring.
  • (ii) A coherent and robust transient eddy response to the EY forcing is observed and serves as a pathway for the downstream surface climate response. The transient response is seen in eddy variance statistics and storm-track diagnostics and propagates far downstream from the forcing region. The response reaches a maximum in amplitude and extent in spring, when it stretches across northern Eurasia and the North Pacific. The nature of the snow-forced signals evolution is indicative of considerable downstream development and is consistent with known storm-track dynamics.
  • (iii) This transient response is a result of the persistent steepened temperature gradients created by the anomalous snow conditions. These gradients contribute to enhanced baroclinicity over the storm-track entrance regions. The sustained nature of the forcing allows the signal to strengthen and propagate downstream, enhancing baroclinicity over areas well outside the traditional storm tracks.

We show that NA snow anomalies can initiate and maintain a physically plausible atmospheric teleconnection with consequences for Eurasian climate, via enhanced North Atlantic storm-track activity. Such a transient eddy response has not previously been reported for a snow forcing over NA, and so it represents a new contribution to our understanding of snow climate interactions.

Acknowledgments

The authors would like to thank Yan Ge for providing quality-checked monthly gridded snow depth and snow frequency data and the two reviewers for their helpful comments and suggestions. Support for this research is provided by NOAA Grant NAO30AR4320179 and NASA ESS Fellowship NNX06AG28H.

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Fig. 1.
Fig. 1.

Annual cycle of area-weighted average NA SWE (solid lines) and snow extent (SNE; dashed lines) for high- and low-snow forcing.

Citation: Journal of Climate 23, 3; 10.1175/2009JCLI3219.1

Fig. 2.
Fig. 2.

Seasonal SWE (black contour lines) and surface temperature (color-filled contours) response to EY snow forcing during (a) fall, (b) winter, (c) spring, and (d) summer. Positive SWE contours (solid lines) drawn at 0.5, 5, 15, and 25 cm. Negative SWE contours (dashed lines) drawn at −0.5, −1, −2, and −3 cm. Only statistically significant (>95%) surface temperatures are plotted.

Citation: Journal of Climate 23, 3; 10.1175/2009JCLI3219.1

Fig. 3.
Fig. 3.

Three-month running mean 500-hPa geopotential height variance ()−1/2 response to EY snow forcing. Seasons begin with fall (SON), and progress, ending with ASO. Contours drawn at ±1, 2, 3, and 4 m; solid (dashed) contours denote a positive (negative) response. 90% (95%) statistical significance indicated with light (dark) gray shading.

Citation: Journal of Climate 23, 3; 10.1175/2009JCLI3219.1

Fig. 4.
Fig. 4.

Spring season transient activity response to EY snow forcing. (a) 500-hPa EKE; contours drawn at ±1, 2, 3, and 4 m2 s−1. (b) 750-hPa meridional temperature flux ; contours drawn at ±0.25°, 0.75°, 1.0°, and 1.5°C m s−1. (c) 750-hPa vertical temperature flux ; contours drawn at 0.01°K Pa s−1 intervals. (d) 750-hPa EGR; contours drawn at ±0.02, 0.04, 0.08, and 0.12 day−1. Solid (dashed) contours denote a positive (negative) response. 90% (95%) statistical significance indicated with light (dark) gray shading.

Citation: Journal of Climate 23, 3; 10.1175/2009JCLI3219.1

Fig. 5.
Fig. 5.

Spring season spatial gradient response to EY snow forcing at 750 hPa. (a) Vertical shear −∂u/∂p; contours drawn at 0.001 m s−1 intervals. (b) Meridional temperature gradient |∂T |/∂y; contours at ±4°, 6°, 8°, 10°, 12°, 16°, 20°, and 24°C 1000 km−1. Solid (dashed) contours denote a positive (negative) response. 90% (95%) statistical significance indicated with light (dark) gray shading.

Citation: Journal of Climate 23, 3; 10.1175/2009JCLI3219.1

Fig. 6.
Fig. 6.

Spring season latitude–pressure profile response to EY snow forcing averaged over NA longitudes (140°–60°W). (a) Zonal wind u; contours drawn at ±0.5, 1, 1.5, and 2 m s−1. (b) Air temperature T; contours drawn at ±0.5°, 1°, 1.5°, and 2°C. Solid (dashed) contours denote a positive (negative) response. 90% (95%) statistical significance indicated with light (dark) gray shading.

Citation: Journal of Climate 23, 3; 10.1175/2009JCLI3219.1

Fig. 7.
Fig. 7.

Column-averaged diabatic heating response to EY forcing during (a) winter, and (b) spring. Contours drawn at ±0.5 K day−1 intervals. Solid (dashed) contours denote a positive (negative) response. 90% (95%) statistical significance indicated with light (dark) gray shading.

Citation: Journal of Climate 23, 3; 10.1175/2009JCLI3219.1

Fig. 8.
Fig. 8.

As in Fig. 2, but for FS snow.

Citation: Journal of Climate 23, 3; 10.1175/2009JCLI3219.1

Fig. 9.
Fig. 9.

As in Fig. 3, but for FS snow.

Citation: Journal of Climate 23, 3; 10.1175/2009JCLI3219.1

Fig. 10.
Fig. 10.

Seasonal 750-hPa meridional temperature gradient |∂T |/∂y response to (a) EY and (b) FS snow forcing. (left) Fall, (middle) winter, and (right) spring seasons. Contours drawn at ±4°, 6°, 8°, 10°, and 12°C 1000 km−1. Solid (dashed) contours denote a positive (negative) response. 90% (95%) statistical significance indicated with light (dark) gray shading.

Citation: Journal of Climate 23, 3; 10.1175/2009JCLI3219.1

Fig. 11.
Fig. 11.

Winter 850-hPa stationary wave streamfunction response to FS forcing. Contours drawn at ±0.5, 0.75, 1.0, and 1.25 × 106 m2 s−1. Solid (dashed) contours denote a positive (negative) response. 90% (95%) statistical significance indicated with light (dark) gray shading.

Citation: Journal of Climate 23, 3; 10.1175/2009JCLI3219.1

Fig. 12.
Fig. 12.

The 500-hPa stationary wave streamfunction ensemble mean (EY low snow) during (a) winter and (c) spring; contours drawn at ±2 × 106 m2 s−1 intervals. 500-hPa stationary wave streamfunction response to EY forcing during (b) winter and (d) spring; contours drawn at ±0.5 × 106 m2 s−1 intervals. Solid (dashed) contours denote a positive (negative) response. 90% (95%) statistical significance indicated with light (dark) gray shading.

Citation: Journal of Climate 23, 3; 10.1175/2009JCLI3219.1

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