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1. Introduction A stratospheric sudden warming (SSW) is a typical manifestation of troposphere–stratosphere interaction in winter (e.g., Charney and Drazin 1961 ; Charlton and Polvani 2007 ). Since the first dynamical model of SSW was established by Matsuno (1971) , subsequent studies have suggested that the process of SSW is closely related to the upward propagation of planetary waves (PWs), especially waves 1 and 2, from the troposphere to the stratosphere (e.g., Nishii et al. 2011
1. Introduction A stratospheric sudden warming (SSW) is a typical manifestation of troposphere–stratosphere interaction in winter (e.g., Charney and Drazin 1961 ; Charlton and Polvani 2007 ). Since the first dynamical model of SSW was established by Matsuno (1971) , subsequent studies have suggested that the process of SSW is closely related to the upward propagation of planetary waves (PWs), especially waves 1 and 2, from the troposphere to the stratosphere (e.g., Nishii et al. 2011
upper levels ( Rivière 2009 ). On the contrary, the planetary, low-frequency waves act to hasten the short-term decay of the zonal wind anomalies during the first week following their peak ( Feldstein and Lee 1998 ; Watterson 2002 ). By analyzing observational datasets, Lorenz and Hartmann (2003) showed that the jet acts as a waveguide for these waves; so they propagate into the jet and remove momentum from it. This general behavior of planetary waves is well reproduced in simple models ( O
upper levels ( Rivière 2009 ). On the contrary, the planetary, low-frequency waves act to hasten the short-term decay of the zonal wind anomalies during the first week following their peak ( Feldstein and Lee 1998 ; Watterson 2002 ). By analyzing observational datasets, Lorenz and Hartmann (2003) showed that the jet acts as a waveguide for these waves; so they propagate into the jet and remove momentum from it. This general behavior of planetary waves is well reproduced in simple models ( O
is eddy dominated, whereas during winter the flow is closer to angular momentum conserving. In a zonal-mean framework, the monsoon–anticyclone system can be considered as a planetary-scale Rossby wave driven by land–ocean (east–west) heating asymmetries [following Gill (1980) ] with associated planetary-scale wave transport. Lorenz (1969, 1984) made a clear distinction between the “ideal Hadley circulation,” which is zonally symmetric and the “modified Hadley circulation,” which includes east
is eddy dominated, whereas during winter the flow is closer to angular momentum conserving. In a zonal-mean framework, the monsoon–anticyclone system can be considered as a planetary-scale Rossby wave driven by land–ocean (east–west) heating asymmetries [following Gill (1980) ] with associated planetary-scale wave transport. Lorenz (1969, 1984) made a clear distinction between the “ideal Hadley circulation,” which is zonally symmetric and the “modified Hadley circulation,” which includes east
( Barnston and Livezey 1987 ). In spite of these efforts by the previous studies, mechanisms for the interannual variability of the East Asian winter monsoon have not been fully clarified yet. The Siberian high and Aleutian low, which constitute the East Asian winter monsoon system, are the prominent semipermanent pressure systems at the surface ( Fig. 1a ), as a surface manifestation of the planetary waves in winter ( Lau 1979 ; Wallace 1983 ). Figures 1b and 1c show the zonally asymmetric component
( Barnston and Livezey 1987 ). In spite of these efforts by the previous studies, mechanisms for the interannual variability of the East Asian winter monsoon have not been fully clarified yet. The Siberian high and Aleutian low, which constitute the East Asian winter monsoon system, are the prominent semipermanent pressure systems at the surface ( Fig. 1a ), as a surface manifestation of the planetary waves in winter ( Lau 1979 ; Wallace 1983 ). Figures 1b and 1c show the zonally asymmetric component
1. Introduction Planetary wave activity accounts for a large portion of the spatial and temporal variability of the stratosphere. Its interaction with the zonal-mean flow affects the strength and the duration of the polar vortex (e.g., Mechoso et al. 1985 ; Polvani and Plumb 1992 ). Planetary wave breaking is the major driver of the equator-to-pole hemispheric Brewer–Dobson circulation (BDC) in the stratosphere (e.g., Rosenlof and Holton 1993 ; Holton et al. 1995 ). Stratospheric planetary
1. Introduction Planetary wave activity accounts for a large portion of the spatial and temporal variability of the stratosphere. Its interaction with the zonal-mean flow affects the strength and the duration of the polar vortex (e.g., Mechoso et al. 1985 ; Polvani and Plumb 1992 ). Planetary wave breaking is the major driver of the equator-to-pole hemispheric Brewer–Dobson circulation (BDC) in the stratosphere (e.g., Rosenlof and Holton 1993 ; Holton et al. 1995 ). Stratospheric planetary
1. Introduction Planetary-scale waves play an important role in determining the climate of the winter hemisphere. They are a particularly important component of the dynamics of the winter season in the Northern Hemisphere where they are readily forced by the combination of orography and continent–ocean heating asymmetries. Planetary waves are well known to be one of the dominant modes of coupling between the stratosphere and troposphere because of their significant vertical propagation. They
1. Introduction Planetary-scale waves play an important role in determining the climate of the winter hemisphere. They are a particularly important component of the dynamics of the winter season in the Northern Hemisphere where they are readily forced by the combination of orography and continent–ocean heating asymmetries. Planetary waves are well known to be one of the dominant modes of coupling between the stratosphere and troposphere because of their significant vertical propagation. They
their ability to transport heat poleward. We find that the poleward sensible heat transport is 1.17 PW throughout the entire column—a value in good agreement with previous studies ( Peixoto and Oort 1992 ; Trenberth and Stepaniak 2003 ). Partitioning the transport between tropospheric (below 300 hPa) and stratospheric contributions, we find transports of 0.46 and 0.71 PW, respectively. Insomuch as there are mechanisms that excite transient planetary-scale waves that act to amplify through
their ability to transport heat poleward. We find that the poleward sensible heat transport is 1.17 PW throughout the entire column—a value in good agreement with previous studies ( Peixoto and Oort 1992 ; Trenberth and Stepaniak 2003 ). Partitioning the transport between tropospheric (below 300 hPa) and stratospheric contributions, we find transports of 0.46 and 0.71 PW, respectively. Insomuch as there are mechanisms that excite transient planetary-scale waves that act to amplify through
et al. 2016 ), and model biases ( Mitchell et al. 2015b ; Dhomse et al. 2016 ). It has been widely accepted that the atmospheric response to the initially small-magnitude solar radiative forcing must involve amplification via nonlinear processes ( Gray et al. 2010 ). One classic mechanism involves the dynamical interaction between upward-propagating planetary-scale Rossby waves (planetary waves hereafter) and the background westerly flow in the winter stratosphere. When a critical layer in a
et al. 2016 ), and model biases ( Mitchell et al. 2015b ; Dhomse et al. 2016 ). It has been widely accepted that the atmospheric response to the initially small-magnitude solar radiative forcing must involve amplification via nonlinear processes ( Gray et al. 2010 ). One classic mechanism involves the dynamical interaction between upward-propagating planetary-scale Rossby waves (planetary waves hereafter) and the background westerly flow in the winter stratosphere. When a critical layer in a
warming do not return to westerly until the final cooling takes place during the fall. The date on which the polar vortex breaks up, the so-called final warming onset time, varies from year to year. The interannual variability in the timing of the final warmings depends on the strength of planetary wave forcing ( Farrara and Mechoso 1986 ), and it can have a large impact on the chemical depletion of stratospheric ozone ( Salby and Callaghan 2007 ). A late final warming is often associated with more
warming do not return to westerly until the final cooling takes place during the fall. The date on which the polar vortex breaks up, the so-called final warming onset time, varies from year to year. The interannual variability in the timing of the final warmings depends on the strength of planetary wave forcing ( Farrara and Mechoso 1986 ), and it can have a large impact on the chemical depletion of stratospheric ozone ( Salby and Callaghan 2007 ). A late final warming is often associated with more
forced planetary-scale waves can warm the Arctic ( Lee et al. 2011a , b ; Yoo et al. 2012 ; Lee 2012 ; Ding et al. 2014 ), we ask the following: Do these planetary-scale waves tap this vast reservoir of ZAPE without relying on the flux–gradient relationship? This was conjectured previously ( Lee 2014 ), but its validity has not been tested. In this study, we address this question by comparing the life cycles of planetary-scale and synoptic-scale waves. Because Rossby waves propagate from the
forced planetary-scale waves can warm the Arctic ( Lee et al. 2011a , b ; Yoo et al. 2012 ; Lee 2012 ; Ding et al. 2014 ), we ask the following: Do these planetary-scale waves tap this vast reservoir of ZAPE without relying on the flux–gradient relationship? This was conjectured previously ( Lee 2014 ), but its validity has not been tested. In this study, we address this question by comparing the life cycles of planetary-scale and synoptic-scale waves. Because Rossby waves propagate from the
