• Baldwin, M. P., and Coauthors, 2001: The quasi-biennial oscillation. Rev. Geophys, 39 , 179229.

  • Crooks, S., , and L. J. Gray, 2004: An examination of the influence of the 11-year solar cycle on stratospheric climate using the ERA-40 data. J. Climate, submitted.

    • Search Google Scholar
    • Export Citation
  • Dunkerton, T. J., , and M. P. Baldwin, 1992: Modes of interannual variability in the stratosphere. Geophys. Res. Lett, 19 , 4952.

  • Gray, L. J., 2003: The influence of the equatorial upper stratosphere on stratospheric sudden warmings. Geophys. Res. Lett.,30, 1166, doi:10.1029/2002GL016430.

    • Search Google Scholar
    • Export Citation
  • Gray, L. J., , E. F. Drysdale, , T. J. Dunkerton, , and B. N. Lawrence, 2001a: Model studies of the interannual variability of the northern hemisphere stratospheric winter circulation: The role of the quasi-biennial oscillation. Quart. J. Roy. Meteor. Soc, 127 , 14131432.

    • Search Google Scholar
    • Export Citation
  • Gray, L. J., , S. J. Phipps, , T. J. Dunkerton, , M. P. Baldwin, , E. F. Drysdale, , and M. R. Allen, 2001b: A data study of the influence of the equatorial upper stratosphere on northern hemisphere stratospheric sudden warmings. Quart. J. Roy. Meteor. Soc, 127 , 19852003.

    • Search Google Scholar
    • Export Citation
  • Gray, L. J., , S. Sparrow, , M. Juckes, , A. O'Neill, , and D. G. Andrews, 2003: Flow regimes in the winter stratosphere of the northern hemisphere. Quart. J. Roy. Meteor. Soc, 129 , 925945.

    • Search Google Scholar
    • Export Citation
  • Haigh, J. D., 2003: The effect of solar variability on the Earth's climate. Philos. Trans. Roy. Soc. London, 361 , 95111.

  • Harvey, V. L., , M. H. Hitchman, , R. B. Pearce, , and T. D. Fairly, 1999: Tropical aerosol in the Aleutian High. J. Geophys. Res, 104 (D6) 62816290.

    • Search Google Scholar
    • Export Citation
  • Haynes, P. H., 1985: Nonlinear instability of a Rossby-wave critical layer. J. Fluid Mech, 161 , 493511.

  • Hitchman, M. H., , and C. B. Leovy, 1986: Evolution of the zonal mean state in the equatorial middle stratosphere during October 1978–May 1979. J. Atmos. Sci, 43 , 31593176.

    • Search Google Scholar
    • Export Citation
  • Holton, J. R., , and H-C. Tan, 1980: The influence of the equatorial quasi-biennial oscillation on the global circulation at 50 mb. J. Atmos. Sci, 37 , 22002208.

    • Search Google Scholar
    • Export Citation
  • Holton, J. R., , and H-C. Tan, 1982: The quasi-biennial oscillation in the Northern Hemisphere lower stratosphere. J. Meteor. Soc. Japan, 60 , 140148.

    • Search Google Scholar
    • Export Citation
  • Hood, L. L., 2004: Effects of solar UV variability on the stratosphere. Solar Variability and Its Effect on the Earth's Atmosphere and Climate System. Geophys. Monogr, No. 141, Amer. Geophys. Union,. 283303.

    • Search Google Scholar
    • Export Citation
  • Hood, L. L., , and B. E. Soukharev, 2003: Quasi-decadal variability of the tropical lower stratosphere: The role of extratropical wave forcing. J. Atmos. Sci, 60 , 23892403.

    • Search Google Scholar
    • Export Citation
  • Killworth, P. D., , and M. E. McIntyre, 1985: Do Rossby-wave critical layers absorb, reflect or overreflect? J. Fluid Mech, 161 , 449492.

    • Search Google Scholar
    • Export Citation
  • Kodera, K., , and Y. Kuroda, 2002: Dynamical response to the solar cycle: Winter stratopause and lower stratosphere. J. Geophys. Res.,107, 4749, doi:10.1029/2002JD002224.

    • Search Google Scholar
    • Export Citation
  • Kodera, K., , Y. Kuroda, , and S. Pawson, 2000: Stratospheric sudden warmings and slowly propagating zonal-mean zonal wind anomalies. J. Geophys. Res, 105 , 1235112359.

    • Search Google Scholar
    • Export Citation
  • Labitzke, K., 1987: Sunspots, the QBO and the stratospheric temperature in the north polar region. Geophys. Res. Lett, 14 , 535537.

  • Labitzke, K., 2003: The global signal of the 11-year sunspot cycle in the atmosphere: When do we need the QBO? Meteor. Z, 12 , 209216.

    • Search Google Scholar
    • Export Citation
  • Labitzke, K., , and H. van Loon, 1988: Association between the 11-year solar cycle, the QBO and the atmosphere. Part I: The troposphere and stratosphere in northern hemisphere in winter. J. Atmos. Terr. Phys, 50 , 197206.

    • Search Google Scholar
    • Export Citation
  • Labitzke, K., , and H. van Loon, 1999: The Stratosphere: Phenomena, History and Relevance. Springer-Verlag, 179 pp.

  • Labitzke, K., , and H. van Loon, 2000: The QBO effect on the global stratosphere in northern winter. J. Atmos. Sol.-Terr. Phys, 62 , 621628.

    • Search Google Scholar
    • Export Citation
  • Lindzen, R. S., , and K. K. Tung, 1978: Wave overreflection and shear instability. J. Atmos. Sci, 35 , 16261632.

  • Matthes, K., , U. Langematz, , L. J. Gray, , K. Kodera, , and K. Labitzke, 2004: Improved 11-year solar signal in the FUB-CMAM. J. Geophys. Res.,109, D06101, doi:10.1029/2003JD004012.

    • Search Google Scholar
    • Export Citation
  • McCormack, J. P., 2003: The influence of the 11-year solar cycle on the quasi-biennial oscillation. Geophys. Res. Lett.,30, 2162, doi:10.1029/2003GL018314.

    • Search Google Scholar
    • Export Citation
  • Naito, Y., , and I. Hirota, 1997: Interannual variability of the northern winter stratospheric circulation related to the QBO and the solar cycle. J. Meteor. Soc. Japan, 75 , 925937.

    • Search Google Scholar
    • Export Citation
  • O'Neill, A., , and V. Pope, 1988: Simulations of linear and nonlinear disturbances in the stratosphere. Quart. J. Roy. Meteor. Soc, 114 , 10631110.

    • Search Google Scholar
    • Export Citation
  • O'Sullivan, D., , and T. J. Dunkerton, 1994: Seasonal development of the extratropical QBO in a numerical model of the middle atmosphere. J. Atmos. Sci, 51 , 37063721.

    • Search Google Scholar
    • Export Citation
  • Polvani, L. M., , and R. Saravanan, 2000: The three dimensional structure of breaking Rossby waves in the polar wintertime stratosphere. J. Atmos. Sci, 57 , 36633685.

    • Search Google Scholar
    • Export Citation
  • Robinson, W. A., 1986: The behavior of planetary wave 2 in preconditioned zonal flows. J. Atmos. Sci, 43 , 31093121.

  • Salby, M., , and P. Callaghan, 2000: Connection between the solar cycle and the QBO: The missing link. J. Climate, 13 , 328338.

  • Scaife, A. A., , and I. N. James, 2000: Response of the stratosphere to interannual variability of tropospheric planetary waves. Quart. J. Roy. Meteor. Soc, 126 , 275297.

    • Search Google Scholar
    • Export Citation
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    (a) Time series of zonally averaged, J–F-averaged NP temperatures (K) at 30 hPa (approximately 24 km) (solid line) and zonally averaged zonal wind (m s−1) at the equator at 44 hPa (dotted line). (b) Time series of 10.7-cm solar radio flux (10−22 W m−2 Hz−1). (c) Time series of zonally averaged, J–F-averaged NP temperatures (K) at 30 hPa (solid line) for years in which the QBO was in its west phase. The 10.7-cm solar flux (10−22 W m−2 Hz−1) is also superposed (dotted lines). (d) As in (c) but for the QBO/E phase years

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    QBO composites: monthly averaged latitude–height sections of zonally averaged zonal wind anomalies (m s−1) for the period 1979–2001 in the months Oct–Apr. The data have been composited into (left) QBO/E and (middle) QBO/W winters, as described in the text. The right-hand column shows the QBO/E minus QBO/W difference. Contour interval is 2 m s−1, dashed contour is zero, and dotted contours are negative values. In the left and middle columns, shaded regions indicate values less than −4 m s−1. In the right-hand column, the light and dark shading indicate statistical confidence at the 75% and 95% levels, respectively, using a standard Student's t test

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    As in Fig. 2 but for SC composites. The data have been composited into (left) Smin and (middle) Smax years, as described in the text. The right-hand column shows the Smax minus Smin difference

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    Time series of area-weighted, zonally averaged modeled temperatures (K) north of 62.5°N at 32 km from (a) the control run, (b) E20–30, (c) E25–35, (d) E30–40, (e) E35–45, and (f) E40–50

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    Ensemble-averaged Eliassen–Palm flux diagnostics for experiments (left) E25–35 and (middle) E30–40 and (right) their difference (E25– 35 minus E30–40) at various days throughout the integration. The contours of E–P flux divergence are in m s−1, according to the “acceleration” scaling described in Gray et al. (2003). Solid contours are positive, the dashed contour is zero, and dotted contours are negative. The shading in the first two columns highlights regions less than −2 m s−1. The shading in the difference plots indicates significance at the 99% level

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    Time series of area-weighted, zonally averaged temperature (K) north of 62.5°N at various heights from a typical ensemble member of the (a) E25–35 experiment and (b) E30–40 experiment

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    Time series of area-weighted, zonally averaged temperature (K) at 44 km at each latitude grid box from (a) E25–35 and (b) E30– 40. The plots span from 2.5° (top line) to 87.5°N (bottom line) in 5° latitude intervals

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    Typical latitude–height sections of zonally averaged zonal winds (m s−1) on day 10 from (top) S40–50 experiment, (middle) the control experiment, and (bottom) the difference between them

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    As in Fig. 4 but from (a) the S40–50 ensemble and (b) the control ensemble

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    Typical distribution of NH Ertel potential vorticity on the 1650-K isentropic surface (10−4 K m−1 s−1) at day 50 from (a) the S40–50 experiment and (b) the control experiment

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    As in Fig. 5 but for (left) experiment S40–50 and (middle) the control experiment and (right) their difference (S40–50 minus control) at various days throughout the integration. The arrows in the top two rows have been magnified by a factor of 20 and 30, respectively, compared with the remaining plots to improve legibility

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    Typical latitude–height sections of scaled Ertel potential vorticity (as described in Polvani and Saravanan 2000) at day 20 from (left) the S40–50 experiment, (middle) the control experiment, and (right) the difference between them (10−4 K m−1 s−1)

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    Monthly averaged latitude–height sections of zonally averaged zonal wind anomalies (m s−1) for the period 1958–2001 in the months Oct–Apr. The data have been composited into (first column) Smin/E, (second column) Smax/W, (third column) Smax/E, and (fourth column) Smin/W years, as described in the text. Contour interval is 5 m s−1 with additional ±2 m s−1 contours also included. Solid contours are positive; dashed contour is zero. Dotted contours are negative values. The light and dark shading indicate statistical confidence at the 75% and 95% levels, respectively. These estimates are for data values different from zero, using a standard Student's t test

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    Time series of daily averaged NP temperature (K) at 50 hPa. The data have been composited into Smin/E, Smin/W, Smax/E and Smax/W years, as described in the text

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Solar and QBO Influences on the Timing of Stratospheric Sudden Warmings

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  • 1 Reading University, Reading, United Kingdom
  • | 2 Oxford University, Oxford, United Kingdom
  • | 3 Rutherford Appleton Laboratory, Chilton, United Kingdom
  • | 4 Oxford University, Oxford, United Kingdom
  • | 5 Rutherford Appleton Laboratory, Chilton, United Kingdom
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Abstract

The interaction of the 11-yr solar cycle (SC) and the quasi-biennial oscillation (QBO) and their influence on the Northern Hemisphere (NH) polar vortex are studied using idealized model experiments and ECMWF Re-Analysis (ERA-40). In the model experiments, the sensitivity of the NH polar vortex to imposed easterlies at equatorial/subtropical latitudes over various height ranges is tested to explore the possible influence from zonal wind anomalies associated with the QBO and the 11-yr SC in those regions. The experiments show that the timing of the modeled stratospheric sudden warmings (SSWs) is sensitive to the imposed easterlies at the equator/subtropics. When easterlies are imposed in the equatorial or subtropical upper stratosphere, the onset of the SSWs is earlier.

A mechanism is proposed in which zonal wind anomalies in the equatorial/subtropical upper stratosphere associated with the QBO and 11-yr SC either reinforce each other or cancel each other out. When they reinforce, as in Smin–QBO-east (Smin/E) and Smax–QBO-west (Smax/W), it is suggested that the resulting anomaly is large enough to influence the development of the Aleutian high and hence the time of onset of the SSWs. Although highly speculative, this mechanism may help to understand the puzzling observations that major warmings often occur in Smax/W years even though there is no strong waveguide provided by the QBO winds in the lower equatorial stratosphere.

The ERA-40 data are used to investigate the QBO and solar signals and to determine whether the observations support the proposed mechanism. Composites of ERA-40 zonally averaged zonal winds based on the QBO (E/W), the SC (min/max), and both (Smin/E, Smin/W, Smax/E, Smax/W) are examined, with emphasis on the Northern Hemisphere winter vortex evolution. The major findings are that QBO/E years are more disturbed than QBO/W years, primarily during early winter. Sudden warmings in Smax years tend to occur later than in Smin years. Midwinter warmings are more likely during Smin/E and Smax/W years, although the latter result is only barely statistically significant at the 75% level. The data show some support for the proposed mechanism, but many more years are required before it can be fully tested.

Corresponding author address: Prof. Lesley Gray, NCAS Centre for Global Atmospheric Modelling, Meteorology Dept., Reading University, Earley Gate, P.O. Box 243, Reading RG6 6BB, United Kingdom. Email: Lesley@met.reading.ac.uk

Abstract

The interaction of the 11-yr solar cycle (SC) and the quasi-biennial oscillation (QBO) and their influence on the Northern Hemisphere (NH) polar vortex are studied using idealized model experiments and ECMWF Re-Analysis (ERA-40). In the model experiments, the sensitivity of the NH polar vortex to imposed easterlies at equatorial/subtropical latitudes over various height ranges is tested to explore the possible influence from zonal wind anomalies associated with the QBO and the 11-yr SC in those regions. The experiments show that the timing of the modeled stratospheric sudden warmings (SSWs) is sensitive to the imposed easterlies at the equator/subtropics. When easterlies are imposed in the equatorial or subtropical upper stratosphere, the onset of the SSWs is earlier.

A mechanism is proposed in which zonal wind anomalies in the equatorial/subtropical upper stratosphere associated with the QBO and 11-yr SC either reinforce each other or cancel each other out. When they reinforce, as in Smin–QBO-east (Smin/E) and Smax–QBO-west (Smax/W), it is suggested that the resulting anomaly is large enough to influence the development of the Aleutian high and hence the time of onset of the SSWs. Although highly speculative, this mechanism may help to understand the puzzling observations that major warmings often occur in Smax/W years even though there is no strong waveguide provided by the QBO winds in the lower equatorial stratosphere.

The ERA-40 data are used to investigate the QBO and solar signals and to determine whether the observations support the proposed mechanism. Composites of ERA-40 zonally averaged zonal winds based on the QBO (E/W), the SC (min/max), and both (Smin/E, Smin/W, Smax/E, Smax/W) are examined, with emphasis on the Northern Hemisphere winter vortex evolution. The major findings are that QBO/E years are more disturbed than QBO/W years, primarily during early winter. Sudden warmings in Smax years tend to occur later than in Smin years. Midwinter warmings are more likely during Smin/E and Smax/W years, although the latter result is only barely statistically significant at the 75% level. The data show some support for the proposed mechanism, but many more years are required before it can be fully tested.

Corresponding author address: Prof. Lesley Gray, NCAS Centre for Global Atmospheric Modelling, Meteorology Dept., Reading University, Earley Gate, P.O. Box 243, Reading RG6 6BB, United Kingdom. Email: Lesley@met.reading.ac.uk

1. Introduction

Interesting correlations between the 11-yr solar cycle (SC) and the quasi-biennial oscillation (QBO) and the frequency of sudden stratospheric warmings (SSWs) have been observed for many years. Correlations between the 10.7-cm radio flux and both temperature and geopotential height in the lower stratosphere were first reported by Labitzke (1987) and Labitzke and van Loon (1988, hereafter LvL88). A correlation between the direction of the QBO in the lower stratosphere and Northern Hemisphere (NH) polar winter temperatures was first reported by Holton and Tan (1980, 1982). They showed that on average, when the equatorial winds near 50 hPa are easterly (westerly), the NH polar winter tends to be warmer and more disturbed (colder and less disturbed). Labitzke (1987) and LvL88 examined this relationship for solar minimum and solar maximum years separately. They confirmed this relationship in solar minimum years but found that it reversed under solar maximum years. Gray et al. (2001b, hereafter G2001b) examined both radiosonde and rocketsonde data near the equator and found a similar result. Using radiosonde data from all years in the period 1955–99 (i.e., four 11-yr SCs) G2001b found a negative correlation of only −0.2 between equatorial winds in the lower stratosphere and January–February average North Pole (NP) temperatures. However, this changed to −0.7 when only solar minimum years were considered. When only solar maximum years were used, the correlation was positive but very weak (+0.1). [For a more recent summary of the QBO see Baldwin et al. (2001) and of the solar– QBO interaction see Labitzke and van Loon (1999, 2000) and Labitzke (2004).]

The mechanism for the interaction of the 11-yr SC and QBO signals is not well understood. The QBO influence from the equatorial lower stratosphere has been explained in terms of the presence of a zero wind line in the northern subtropics in QBO-east phase (QBO/E) years (i.e., years when the zonal flow is easterly). This confines the propagation of planetary-scale waves closer to the pole than in QBO-west phase (QBO/W) years, when the zero-wind line is in the southern subtropics. However, the primary influence from the 11-yr SC is thought to be on the temperatures in the equatorial upper stratosphere (e.g., Hood 2004 and references therein). One of the greatest challenges in this research area is to find a plausible mechanism for the extension of a significant solar cycle influence into the lower stratosphere.

Observations indicate a small solar signal in equatorial lower-stratospheric temperatures (Hood 2004; Haigh 2003; Crooks and Gray 2004, manuscript submitted to J. Climate) and a solar modulation of the length of the west phase of the QBO in the lower stratosphere (Salby and Callaghan 2000). It is not clear whether this is a direct response to changes in solar irradiance or whether it is an indirect response either via an “equatorial route,” for example, a modulation of the rate of descent of the QBO at the equator (see, e.g., McCormack 2003) or via a “polar route,” for example, a modulation of the polar vortex that feeds back to lower latitudes via the meridional circulation (e.g., Kodera and Kuroda 2002; Hood and Soukharev 2003). Similarly, it is not clear whether the observed solar signal in the polar lower stratosphere is a result of the solar modulation of the QBO in the lower stratosphere (as observed by Salby and Callaghan 2000) or whether there is some influence from the upper equatorial stratosphere, which then influences the polar vortex at all heights.

Rocketsone data (G2001b) and the European Centre for Medium-Range Weather Forecasts (ECMWF) Re-Analysis (ERA-40) data (Pascoe et al. 2004, manuscript submitted to J. Geophys. Res.) indicate that the QBO extends to at least 60 km, although the amplitude at these higher levels (±5–10 m s−1) is much weaker than at 20–30 km (±20–30 m s−1). Gray (2003) investigated the relative sensitivity of SSWs to the equatorial wind direction in the upper stratosphere (>40 km) compared with the lower stratosphere (<40 km) and found that the sensitivity to the lower-stratospheric equatorial winds was confined to early winter conditions, while the sensitivity to the upper-stratospheric equatorial winds was greater and primarily in mid- and late winter. This suggests that greater sensitivity may compensate for the smaller QBO amplitude in the upper stratosphere. It also allows the possibility of an interaction of the solar and QBO signals in the upper equatorial stratosphere where the 11-yr SC signal has its maximum amplitude and the QBO has a small but not insignificant amplitude (Gray et al. 2001a).

The GCM study of Matthes et al. (2004) has achieved a simulation of the polar SC–QBO interaction by imposing total solar irradiance (TSI) and associated ozone changes at levels appropriate to the observed ranges of the 11-yr SC and by using a simple scheme in which the equatorial zonal winds were relaxed toward typical QBO distributions. They found that, when they relaxed toward a QBO anomaly only in the lower stratosphere (using radiosonde observations to 10 hPa), they did not achieve a realistic SC–QBO polar interaction. However, when the relaxation scheme was extended so that it was employed throughout the whole depth of the equatorial stratosphere (using rocketsonde observations to 1 hPa), they achieved a realistic SC–QBO interaction in the NH winter. Their study therefore suggests that an influence from the equatorial upper stratosphere on the polar vortex is important, supporting the conclusions of Gray et al. (2001a, 2003), who used a more idealized model. It also supports the possibility of the polar route for solar influence to reach the lower stratosphere since their equatorial QBO wind profile was fixed.

In this paper, we build upon previous modeling and data analysis by Gray and collaborators to investigate the possibility that the polar route of SC influence via a modulation of the polar vortex may be identical to the mechanism of QBO influence on the polar vortex. If this is the case, it may help to understand the mechanism for the observed SC–QBO interaction. We use data from the recent ERA-40 dataset to examine the QBO and 11-yr SC signals in the NH winter (section 2). In section 3 we describe a series of idealized model experiments that extend the study of Gray (2003) to investigate the sensitivity of SSWs to imposed equatorial easterlies in selected 10-km-height regions between 20 and 55 km, as suggested by the QBO data analysis, and also to imposed easterlies at the subtropical stratopause region, as suggested by the SC data analysis. Taking into account the results of both the ERA-40 analysis and model experiments, a mechanism is proposed (section 4) that may explain the observed SC-QBO influence on the NH vortex, including the puzzling evidence (LvL88) that under solar maximum conditions major SSWs are more likely to occur when the lower-stratospheric QBO is in its westerly phase, a result contrary to the Holton–Tan mechanism.

Our mechanism involves a sensitivity of SSWs to relatively small wind anomalies in the equatorial and subtropical upper stratosphere. Specifically, an easterly anomaly helps to establish the Aleutian high, which speeds up the onset of the SSW, while a westerly anomaly slows it down. We propose that the superposition of easterly anomalies associated with solar minimum (Smin) conditions and QBO/E conditions results in SSWs in Smin/E years that are consistently earlier than average. Similarly, westerly anomalies associated with Smax conditions and QBO/W years superpose in Smax/W years and result in SSWs that are consistently later than average. In the other combinations, that is, Smin/W and Smax/E the anomalies associated with the SC and QBO are of opposite sign and partially cancel, so there is no consistent speeding up or slowing down of the SSW in those years and hence no signal of SC–QBO interaction emerges.

2. ERA-40 analysis: The QBO and the 11-yr solar cycle

A new synthesis of in situ and remotely sensed measurements for the period since mid-1957 has been employed in this study. It was prepared by ECMWF using their variational data assimilation system and is referred to as the ERA-40 dataset (available online at http:// www.ecmwf.int/research/era). Apart from the extended length of the ERA-40 dataset compared with the earlier ERA-15 dataset, an important difference is that the ERA-40 dataset extends to 0.1 hPa compared with ERA-15 that extended only to 10 hPa. Also, ERA-40 assimilates the Television Infrared Observation Satellite (TIROS) Operational Vertical Sounder (TOVS) radiances directly, while in ERA-15 retrievals of temperature and humidity were assimilated.

In addition to the traditional data used in the assimilation, such as radiosonde observations, Vertical Temperature Profile Radiometer (VTPR), TOVS, and Cloud Motion Winds (CMWs), satellite observations were used in the period 1972–88, and TOVS, Sensor Microwave Imager (SSM/I), European Remote Sensing Satellite (ERS), Advanced TIROS Operational Vertical Sounder (ATOVS) and CMW satellite observations were used in the period 1987–2001. Post-1979 TOVS consisted of the Stratospheric Sounding Unit (SSU)/ High Resolution Infrared Radiation Sounder (HIRS)/ Microwave Sounding Unit (MSU) plus the Advanced Microwave Sounding Unit (AMSU)/HIRS from 1998 onward. The SSU and AMSU observations are predominant at upper-stratospheric levels.

The ERA-40 dataset consists of 6-hourly analyses throughout the period 1957–2001. The data are of high spatial resolution, with a grid spacing close to 125 km in the horizontal (T159) and with 60 levels in the vertical located between the surface and 0.1 hPa (approximately 65 km). The figures shown in this paper are from the 6-hourly analyses on standard pressure surface up to 100 hPa and on model levels between 100 and 0.1 hPa. Caution is required in the interpretation of data above 1 hPa (approximately 50 km) since this will be contaminated by the close proximity of the top of the model.

a. North Polar latitudes

Figure 1 shows the time series over four 11-yr solar cycles (1959–2001) of January–February (J–F) averaged NP temperature anomaly at 30 hPa from the ERA-40 dataset together with the equatorial wind anomaly at 44 hPa in December–January (D–J) of the same winter. This figure may be compared with the corresponding figure from the Berlin stratospheric analyses and equatorial rocketsonde dataset shown in Fig. 2a of G2001b. Both time series show semiregular ∼2-yr variations of alternating sign. There are occasional years in which the variation “skips” a year and the anomaly remains the same sign for two or more consecutive years, for example, 1964–65. The two signals are anticorrelated for several periods of time, for example, 1960–68, 1971–79, 1984–88, 1991–96, and 1999–2001 but are either unrelated or positively correlated during the intervening periods.

As noted by LvL88, Naito and Hirota (1997), and G2001b, the periods of anticorrelation tend to coincide with periods of solar minimum. This is confirmed by comparison of Fig. 1a with Fig. 1b, which shows the time series of 10.7-cm solar radio flux. This relationship is also highlighted in Figs. 1c and 1d, which show the same polar temperature data separated into QBO/E and QBO/W years and are plotted in a similar manner to LvL88 (see also G2001b, their Fig. 4). Generally, when only the QBO/W years are plotted, there is a positive correlation with the 11-yr variation in solar flux and vice versa in QBO/E years. However, Figs. 1c and 1d show that this is by no means always the case. The 11-yr SC and QBO influences are only two of several factors that influence the variability of SSWs. Other possible influences include planetary wave variations associated with ENSO and the inherent nonlinearity of the atmospheric flow. In addition, we note that using J–F-averaged polar temperatures is not necessarily the ideal indicator of SSWs since there is often a period following the warming when the vortex is reestablished and is anomalously cold, so averaging over several months can mask the real signal.

A standard Pearson's correlation of the two time series in Figs. 1c and 1d gave correlations of +0.34 in the QBO/W phase and −0.52 in the QBO/E phase. In order to compare with the correlations shown in Labitzke and van Loon (2000, their Fig. 7) the calculations were repeated for the period 1959–97 using February data only (instead of J–F). The correlation amplitudes increased to +0.59 and −0.51, respectively. Labitzke and van Loon (2000) show maps of correlations with values of the order of +0.7 and −0.1 for the QBO/W and QBO/E phase respectively. Their corresponding values for geopotential fields are +0.7 and −0.4, respectively. Hence, our results agree very well with Labitzke and van Loon (2000) both in terms of the sign of the correlations in each QBO phase and the comparable amplitudes of the correlations. Interestingly, our QBO/W correlations increased slightly, if a small time lag was introduced, so that the 10.7-cm solar flux led the polar temperatures by 4 months, for example, the February correlations then become +0.64 and −0.4, respectively.

b. QBO composites

Figure 2 shows composites of monthly averaged latitude–height cross sections of zonally averaged zonal wind anomalies from the 1979–2001 mean for the months October through April. We use only the data since 1979 because of the improvements to the upper atmosphere representation after that date due to satellite data assimilation. The years have been composited into QBO/E (7 yr) and QBO/W (9 yr) using the D–J-averaged equatorial winds at 44 hPa (approximately 24 km). Years when the amplitude was less than ±5 m s−1 were excluded (see Fig. 1a). The QBO/E minus QBO/W differences are also shown (right-hand column) with shading indicating the 95% and 99% confidence levels. As expected, the equatorial regions are dominated by the QBO, with an anomaly maximum centred at 20–30 km and a second anomaly of opposite sign directly above it centred at 30–40 km. These anomalies descend slowly throughout the period as the phase of the QBO descends.

There is also a third anomaly directly above this, centred at 40–55 km. It has the same sign as the lower-stratospheric anomaly at 20–30 km and thus forms a region of three heightwise layers of alternating signs between 20 and 55 km (see also Pascoe et al. 2004). The amplitude of the third anomaly above 40 km is of the order ±5–10 m s−1 and is significant at the 95% level. This threefold QBO structure can be explained in the following way. Height–time series of zonal winds at the equator indicates that the QBO signal extends upward to at least 50–60 km and the rate of descent of the QBO is slow enough that the next incoming phase is usually present at ∼50 km while an earlier phase of the same sign is still present at 20–30 km, thus forming a threefold structure in height (G2001b). Previous assumptions that the QBO had only a twofold structure over this height region (e.g., Baldwin et al. 2001) were based on analyses that were less accurate than the rocketsonde measurements examined by G2001b or the ERA-40 data shown in Fig. 2. We also note here that this third, upper-level anomaly is unlikely to be caused by anomalous horizontal advection of summer easterlies by the mean circulation associated with SSWs since it is already present in autumn/early winter before the SSWs develop.

At high latitudes, the anomalies in Fig. 2 are as expected from the Holton–Tan relationship. The QBO/E composite (left column) shows an easterly anomaly at high NH latitudes already present by November that gradually increases in magnitude through January. However, the differences between QBO/E and QBO/W composites (right column) are barely significant at the 75% confidence level except in early winter (November). Increasing the number of years of data by using the whole dataset (not shown) does not improve the statistical significance in this region.

c. Eleven-year solar cycle composites

Figure 3 shows monthly averaged latitude–height cross sections of zonally averaged zonal wind anomalies from the 1979–2001 mean for the months October through April. The years have been composited into Smax (11 yr) and Smin (9 yr) phases using the D–J-averaged 10.7-cm radio flux. Years in which the value fell within the range 105–145 (× 10−22) W m−2 Hz−1 were defined as transition years and are excluded from the composites (see Fig. 1b). The Smin composite (left column) has an easterly anomaly above 50 km at the equator in October that moves poleward and downward reaching a maximum amplitude in January. This is evidence of a disturbed vortex in early to midwinter. The Smax composite has a similar anomaly but of the opposite sign. Hence, the Smax minus Smin difference plot (right column) has a strong positive region that moves poleward and downward with time, indicating a colder more stable vortex in Smax years. The confidence levels in December and January for this feature are greater than 95%. These results are very similar to those reported by Kodera et al. (2000).

In February and March, however, this situation is reversed. The Smin composite has a westerly anomaly at high latitudes, evidence of a strong stable vortex, while the Smax composite has an easterly anomaly. The pattern viewed as a whole throughout the winter suggests that Smin winters are characterized by early to midwinter warmings, while Smax winters are characterized by late-winter warmings.

3. Model experiments

The model employed in this study is the U.K. Meteorological Office (Met Office) Stratosphere Mesosphere Model (SMM), identical to that employed by Gray et al. (2003) and Gray (2003). It is a global three-dimensional primitive equation model of the middle atmosphere with horizontal resolution 5° latitude, 5° longitude, and 32 vertical levels equally spaced in log pressure, giving approximately 2-km vertical resolution throughout the model domain 16–80 km. It is a mechanistic model with temperature and horizontal winds as prognostic variables and a lower boundary located near the tropopause at 16 km. A Rayleigh friction scheme is employed above 50 km to simulate the effects of gravity wave breaking in that region.

The design of the experiments is very similar to Gray et al. (2003) and Gray (2003). The model was initialized with NH summertime temperature and wind distributions and then run under perpetual January radiative conditions so that the model established a NH polar vortex (albeit rather more quickly than in reality) in response to radiative cooling. Each model integration was 300 days long. Each experiment consisted of an ensemble of 20 integrations with slightly varying initial conditions. The lower boundary forcing consisted of an imposed zonally symmetric geopotential height field plus a zonally asymmetric wavenumber-one forcing. The peak amplitude of the wave was 250 m at 60°N in all experiments (see Gray et al. 2003 for more details). The forcing was turned on gradually over the first 10 days of the integration so that, by the time the NH background flow was westerly and planetary waves were able to propagate, the lower boundary forcing was fully switched on.

In all experiments an additional forcing was imposed either in equatorial or subtropical latitudes using a Rayleigh-friction-like relaxation toward a target wind value over a specified height and latitude range. A relaxation time scale of 5 days was employed in height regions below 26 km and this was reduced smoothly to 1 day above 32 km to take into account the faster radiative time scale in the upper stratosphere. A latitudinal profile was imposed so that the relaxation time scale was linearly increased to four times the value by ±17.5° either side of the central forcing latitude, that is, to 20 days below 26 km and 4 days above 32 km. The exact height and latitude ranges over which this relaxation mechanism was employed and the target winds employed in each experiment are described in the relevant section.

a. Equatorial forcing

The aim of the first set of experiments was to test the sensitivity of the modeled NP temperature evolution to imposed equatorial easterlies at different heights. The control for this experiment was taken from Gray et al. (2003) in which the equatorial winds were relaxed toward 0 m s−1 at all heights in the range 16–60 km. The evolution of the area-weighted temperature north of 62.5°N at 32 km of this control is shown in Fig. 4a. Since the initial conditions were from typical August distributions, the initial NH zonal flow was easterly. There is a rapid readjustment to the imposed perpetual January radiative conditions in the first 30–40 days with strong polar cooling and the onset of a westerly vortex. A large disturbance is apparent in all ensemble members between days 50 and 90, as evident from the rise and fall of polar temperatures during this period. After day 90 approximately half of the ensemble members proceed to develop a stratospheric warming, peaking at various times between days 170 and 230, while half continue to cool and remain relatively undisturbed throughout the rest of the integration. Thus there is a large degree of variability in this control experiment.

Five sensitivity experiments were then carried out, prompted by the pancakelike structure of the QBO composite anomalies in Fig. 2. The target equatorial wind height profile was modified so that the target equatorial winds were −20 m s−1 in a specified 10-km height region while the remainder of the target profile was kept at 0 m s−1. The transition between −20 and 0 m s−1 at the top and bottom of the 10-km region was carried out smoothly over several grid boxes so that the height region of easterlies actually extends to ∼14 km. The specified 10-km height regions for the five experiments were 20–30, 25–35, 30–40, 35–45, and 40–50 km. We denote these as E20–30, E25–35, E30–40, E35–45, and E40– 50, respectively. This experiment is similar in nature to the experiment described by Gray (2003) in which the equatorial atmosphere was divided rather more crudely into a lower stratosphere (20–40 km) and an upper stratosphere (40–60 km).

Figures 4b–f show the resulting NP temperature evolution from each of the experiments. It is immediately apparent that an imposed easterly band at any height serves to make the winter evolution more disturbed since every ensemble member of every experiment now displays a warming at some stage during the integration, compared with the control experiment in which only about half were disturbed. This result is in good agreement with the theory proposed by Holton and Tan (1982) that equatorial easterlies confine the propagation of planetary waves to high latitudes and result in a more disturbed vortex.

b. Lower versus middle stratosphere

We now consider in more detail the results of each separate experiment. With imposed equatorial easterlies at 20–30 km (E20–30) all of the ensemble members are disturbed, with some peaking earlier than in the control run, at around day 130, while others peak at the same time as in the control, at around day 170–200. When the imposed band of easterlies is raised to 25–35 km (E25–35), the SSWs tend to occur even earlier, with most members now achieving a warming by day 130 and the variability between ensemble members is reduced. As a measure of comparison between the experiments, we note that, in E20–30, 7 of the 20 ensemblers members peaked by day 130 while, in E25–35, 13 of the 20 peaked by day 130.

When the imposed band of easterlies is moved up to 30–40 km (E30–40), two interesting features are apparent. First, the variability is reduced substantially, with almost all of the members following a very similar evolution for at least 200 days. Second, the SSWs occur consistently later than in the previous two experiments, peaking at around day 170. Because of this delay, none of the E30–40 ensemble members peak by day 130.

To further explore this interesting behavior, we show in Fig. 5 a comparison of the Eliassen–Palm (E–P) diagnostics from E25–35 and E30–40 at selected times during the experiment. Because of the relatively small variability between ensemble members in these two experiments, the E–P flux diagnostics are presented as ensemble averages so that an estimate of statistical significance can be achieved. The zero-wind line is also shown as an aid to interpretation. The patterns of the E–P flux divergences from the two experiments are very similar until around day 60–70 when significant differences of a few meters per square second are evident, centered around 45°N, 50 km. By days 90–100 and 100– 110, just preceding the SSWs at day 130 in E25–35, these differences are accentuated.

Finally, at days 120–130 at the peak of the SSWs in E25–35 the zero-wind line in the Tropics extends toward midlatitudes and a zero-wind line is also present at upper levels over the pole in E25–35, as the winter vortex reverses in sign and becomes easterly. This latter zero-wind line penetrates downward over the next 10–20 days as the warming reaches its mature stage. In contrast, E30– 40 shows a similar evolution but delayed by 30–40 days. We note that, while the position of the zero-wind line in the tropical region clearly shows the different imposed equatorial distributions, these differences are really rather small, especially in the subtropics, and neither the E–P flux diagnostics nor the position of the zero-wind line in the lower equatorial/subtropical winter stratosphere provides an obvious explanation for the 30–40-day delay in evolution of the SSWs in E30–40.

In Fig. 6 the evolutions of the polar temperatures at different heights are shown for E25–35 and E30–40. For this comparison, we show results from a single typical ensemble member from each experiment. At 50 km the evolution is similar in the two experiments with rising temperatures from around day 15 when the flow has become sufficiently westerly for the first planetary waves to propagate. This rise in temperature is also seen at the lower levels but at a later time. There are many similarities in the detailed behavior of the two experiments during the lead up to the first warming event except that the events in E30–40 occur slightly later and then evolve, more slowly. One very interesting feature is that although the SSWs in E30–40 take longer to evolve, the actual temperatures achieved at 24 km are 7°–8° warmer at the peak of the warming (223 K at day 175) compared with E25–35 (215 K at day 140). This difference, which is an indication of the extent to which the lower-level vortex has been disrupted, may have important implications for stratospheric influence on the tropospheric circulation below.

In Fig. 7, we further explore the temperature evolution at 44 km by showing the evolution at all latitudes between the equator and the pole. Again, while the overall evolution is rather similar in the two experiments, there are interesting differences. At the equator/subtropics the temperature gradient in E30–40 is greater, an indication that the region of imposed equatorial easterly winds is just below the 44-km level and this level is in a region of positive vertical wind shear as the equatorial wind profile changes from −20 m s−1 at 40 km to 0 m s−1 by around 48 km (see position of the zero-wind line in Fig. 5). In contrast, E25–35 has virtually no temperature gradient between the equator and 20°N since the region of imposed easterlies is well below 44 km in this experiment and there is no equatorial vertical wind shear present. At higher latitudes there are a number of fairly regular small-scale fluctuations, especially in E30–40 between 30° and 80°N from around day 90 to 100. These disturbances were noted by Scaife and James (2000) and Gray et al. (2003) to be an indication of the presence of small anticyclones resulting from the incursion of a tongue of equatorial [low potential vorticity (PV)] air into the subtropics that are then advected around the edge of the vortex and merge with the Aleutian high [see also earlier studies of these anticyclones by O'Neill and Pope (1988) and Harvey et al. (1999)]. These fluctuations are also present in E25– 35 except that at around day 100 there is one large fluctuation that leads to much higher temperatures, forming the start of the warming.

In summary, the earlier onset of the warming in E25– 35 appears to be a result of a single effective event at around day 100, while E30–40 exhibits three to four smaller events between days 95 and 115 before it, too, achieves a dramatic rise in temperatures to form the warming. Examination of the corresponding latitudinal temperature evolutions at lower heights (not shown) also indicates that a similar behavior occurs there, with E30– 40 exhibiting small-scale fluctuations for a longer time before the sudden warming eventually occurs.

c. Lower versus upper stratosphere

As the region of imposed easterlies is moved higher from E30–40 through E35–45 to E40–50 (Fig. 4), the timing of the SSWs becomes more variable and they occur earlier. Indeed, the same proportion of the ensemble achieves a warming by day 130 in E40–50 as in E25–35 and several of the SSWs in E40–50 peak as much as 30 days earlier than in E25–35. This result is consistent with the study of Gray (2003) and related earlier work that found a marked sensitivity of the timing of sudden SSWs to the equatorial winds in the upper stratosphere.

Rather than proceed immediately to a full analysis and comparison of these latter results, which would necessitate the analysis of just a single ensemble member because of the variability in timing of the SSWs in E40– 50, we choose to proceed to the next set of (subtropical) experiments since they are very similar in nature and allow improved estimates of statistical significance.

d. Summary of equatorial forcing experiments

In summary, we have shown that the presence of easterlies at the equator at any height will result in a more disturbed vortex. However, the timing of the resulting SSWs is sensitive to the height of those easterlies. Easterlies in the lower equatorial stratosphere (20–30 km, 25–35 km) tend to speed up the onset of SSWs. Easterlies in the mid equatorial stratosphere (30– 40 km) tend to slow down the onset of SSWs. Easterlies in the upper equatorial stratosphere (35–45 km, 40–50 km) tend to speed up the onset of SSWs. It is not clear how such a small difference in the height range of imposed easterlies at the equator influences the timing of the warming. A full PV budget study of the developing SSWs might help to identify the mechanism of this influence, but is beyond the scope of this study.

e. Subtropical forcing

Leading on from the apparent sensitivity of the polar flow to imposed easterlies in the upper equatorial stratosphere, a similar experiment was conducted but this time relaxing toward 20 m s−1 easterlies in the NH subtropical upper stratosphere instead of the equatorial upper stratosphere. This “subtropical forcing” experiment was prompted by Fig. 3 and previous work, for example, by Kodera et al. (2000), in which an easterly anomaly in Smin is observed to develop in the northern subtropics near the stratospause level in early winter (see November/December, left column in Fig. 3), which then moves poleward and downward as the Aleutian high develops and forces a polar warming event.

In this experiment, which we denote S40–50, an identical zonal-wind relaxation method toward a target wind profile was employed except that the central latitude of the relaxation forcing was at 12.5°N instead of the equator. The easterly forcing therefore extended over the approximate latitude range 0°–30°N. The relaxation was applied only in the height region 40–50 km, and the target winds were −20 m s−1. This forcing was applied only for the first 60 days of the integration to ensure that the forcing did not interfere with the evolution of the sudden warming. Once the warming commences and the vortex evolves toward the characteristic “comma” shape, a substantial tongue of westerlies extends out into the subtropics and any easterly forcing at this stage would therefore interfere with the evolution of the warming itself.

As a “control” for this experiment we used a 20-member ensemble of integrations in which no imposed forcing has been applied anywhere. The zonally averaged zonal-wind fields at day 10 are shown in Fig. 8 from a typical ensemble member from the control and S40–50 experiment to illustrate the effect of the imposed subtropical forcing in S40–50. As with all models with no artificial forcing at the equator (e.g., via our equatorial relaxation method or via a gravity wave drag parameterization scheme), the equatorial regions in the control experiment are easterly everywhere, apart from a small region in the very lowest stratosphere where it takes slightly longer to move away from the initial westerly conditions because the radiative relaxation time is long there. The main effect of the imposed subtropical easterly forcing (see Fig. 8a) is above 40 km: the developing westerly NH polar jet is confined poleward of 30°N instead of spreading across the subtropics into equatorial latitudes. We note that the S40–50 vortex is similar to the “preconditioned” state, long identified as being a precursor to stratospheric warmings, in which the core of the jet becomes more upright and strengthens slightly.

The polar temperature evolution at 32 km of S40–50 and its control experiment is shown in Fig. 9. The control ensemble displays substantial variability, although all members achieve a warming event by day 200, as one would expect knowing that the equatorial winds are weakly easterly at nearly all heights. The S40–50 distribution is quite different. It shows remarkably little variability, with a very early SSW in all ensemble members, peaking by day 100 and in some cases by day 85. If this is compared with the similar equatorial forcing experiment in Fig. 4f, it may be seen that easterly forcing in the subtropics of the upper stratosphere is even more effective at forcing a consistently early SSW than when the easterly forcing is at the equator. This early establishment of the SSWs in S40–50 is seen in Fig. 10, which shows the 1650 K PV distributions at day 50 from a typical S40–50 ensemble member, compared with the corresponding control integration. The S40–50 distribution shows a very well established Aleutian high by day 50 whereas the corresponding control integration has only a weak feature. It is not difficult to understand these differences since the easterly forcing at 0°–30°N in S40–50 is in precisely the latitudinal region to reinforce the anticyclonicity of the Aleutian high and also the travelling anticyclones that are advected around the edge of the vortex and merge with the Aleutian high (see Fig. 10a).

In Fig. 11 we show the E–P flux diagnostics from S40–50 at various stages leading up to the warming. (Note that the scaling of the E–P flux vectors at days 20–30 and 30–40 has been increased for clarity). The main feature at days 20–30 and 30–40 is the dipole structure in E–P flux divergence centred around 25°– 45°N, 50 km. This region coincides with strong horizontal shear in the zonal wind field (see Fig. 8) and indicates horizontal shear instability. Examination of the PV (q) distribution in S40–50 at day 50 in Fig. 10 suggests the presence of a region of negative qy in the vicinity of 30°N and this is confirmed in Fig. 12, which shows the PV cross sections (scaled as in Polvani and Saravanan 2000) on day 20, corresponding to approximately the same time as the E–P flux diagnostics in Fig. 11 (top). Note the minimum in q near 40°N and hence negative qy in the latitude range 25°–40°N and positive qy poleward of this. The dipole structure in the S40–50 E–P flux divergence in Fig. 11 at days 20–40 reflects this change in qy. We note also that the zero-wind line is embedded within the region of negative qy.

The region of EP flux divergence near 30°N, 50 km indicates that wave activity is being focused (overreflected) back into higher latitudes. This is an almost identical situation to that discussed by O'Neill and Pope (1988, their Fig. 5). More wave activity returns to high latitudes than leaves it [for further discussion of this behavior see Lindzen and Tung (1978), Killworth and McIntyre (1985), Haynes (1985), and Robinson (1986)]. In our experiment, the critical layer in the region of 30°–40°N appears to act as a reflector in the early stages of the integration (days 20–40). From around day 40, it becomes more strongly absorbing and the evolution of the E–P flux diagnostics leading up to the warming event become more typical. These diagnostics show that the easterly forcing in the subtropics in the early stages of this experiment has helped to speed up the evolution of the SSWs.

4. Summary and discussion

In summary, the model experiments have shown that the timing of warming events is sensitive to forced easterlies at the equator and in the subtropical upper stratosphere. Imposing easterlies at any height at the equator resulted in earlier SSWs than in the control. The optimum forcing region for producing very early SSWs was in the subtropical upper stratosphere (experiment S40– 50), in which the easterly forcing displaced the zero-wind line into the subtropics above 40 km, in the vicinity of a region of negative qy. Eliassen–Palm flux diagnostics identified that wave activity was focused back into higher latitudes, which resulted in the earlier SSWs. Synoptically, the subtropical easterly forcing enhanced the growth of the anticyclonic Aleutian high so that it quickly reached sufficient amplitude to influence the strength and position of the polar vortex. When the easterly forcing was applied in the upper stratosphere at equatorial latitudes instead of in the subtropics, a similar result was obtained except that the equatorial forcing was not quite so effective as the subtropical forcing so that the SSWs developed a little more slowly and there was more variability between the ensemble members.

Forcing easterlies in the lower stratosphere also produced interesting results suggesting that the timing and depth of penetration of the SSWs was very sensitive to the height of the imposed easterlies. The warmings took approximately 30 days longer to develop when the equatorial easterlies were imposed between 30 and 40 km than when they were imposed between 25 and 35 km. [We note that this latter result is consistent with the results of O'Sullivan and Dunkerton (1994).]

The apparent influence of the equatorial/subtropical wind anomalies above 40 km on the timing of onset of SSWs is very interesting when interpreted in conjunction with the ERA-40 data analysis of section 2. Both QBO composites (Fig. 2) and 11-yr SC composites (Fig. 3) showed statistically significant anomalies in approximately these regions. The amplitude of both the QBO and 11-yr SC zonal wind anomalies above 40 km is around ±5–10 m s−1. The QBO/E and Smin composites tend to show an easterly anomaly in these regions, while the QBO/W and Smax composites show a westerly anomaly. In Smin/E years we may therefore expect the two anomalies to combine to give a relatively strong easterly anomaly. In Smax/W years, we may expect the two anomalies to combine to give a relatively strong westerly anomaly. On the other hand, in Smin/W and Smax/E years we may expect the two anomalies to partially cancel and the resulting wind anomaly to be negligible. This is summarized in Table 1. We can also use the results of the model experiments to predict what effect this may have on the timing of the warmings. In Smin/E years, with a strong easterly anomaly, we may expect to see earlier warmings than in Smax/W years which have a strong westerly anomaly. In the cases when the two anomalies are of opposite sign (Smin/W, Smax/E) and therefore, partially cancel, we may not see any preferred timing since in those cases there is no strong influence from the equatorial/subtropical upper stratosphere.

a. ERA-40 analysis: Solar cycle/QBO composites

We may check the predictions in Table 1 using the ERA-40 dataset. Figure 13 shows composites of monthly averaged latitude–height cross sections of zonally averaged zonal-wind anomalies from the 1958/59– 2000/01 mean for the months October through April. The composites have been compiled on the basis of both the 11-yr solar cycle phase and the QBO phase, giving four plots: Smin/E (7 yr), Smin/W (8 yr), Smax/E (8 yr), Smax/W (8 yr). The criteria employed for defining the QBO and SC phase were as described in section 2. We have used data from all the available years in order to maximize the number of years in each composite. This means that distributions below 10 hPa (∼32 km) are reliable, but above this level there were very few assimilated observations pre-1979 and these must therefore be treated with caution.

The equatorial/subtropical NH anomalies above 40 km in Fig. 13 show a distinct pattern that partially agrees with the predictions summarized in Table 1. In Smin/E and Smax/W there are indeed easterly and westerly anomalies, respectively, in this region and most are statistically significant at the 75% level or higher. However, there is little evidence that they are larger in amplitude than in the Smin/W or Smax/E years, so we do not see a clear signature of the signals reinforcing or canceling out. This is perhaps not surprising, given the paucity of observations that have contributed to the dataset above 32 km. An examination of the corresponding composites using only data since 1979 (not shown), when more satellite data became available, supports the predicted reinforcement/cancelation of anomalies more convincingly, but there are so few years in each composite (average of only 3 yr in each) that the statistical significance of the results is very low.

At NH polar latitudes there are substantial variations in Fig. 13 between the composites. The timing of the NH warmings, indicated by an easterly anomaly at high NH latitudes, is described in this section and the results are summarized in Table 2. In the Smin/E composite (first column) an easterly anomaly builds up at high latitudes in December, reaching maximum amplitude in January, and then moves downwards and weakens in February. In the Smax/E composite (third column) there is an easterly anomaly at high latitudes in December, similar to the anomaly in the Smin/E composite, but this is not perpetuated into the following months. Thus, in both QBO/E composites the early winter is relatively disturbed. This can be compared with, for example, December in the two QBO/W composites (second and fourth columns), both of which have a westerly anomaly at the NP, indicating a stronger, less-disturbed vortex than average. This difference in the early winter behavior between the QBO/E and QBO/W composites is in good agreement with the Holton–Tan relationship. However, this east–west difference is only present in early winter (November/December) and does not continue into midwinter, suggesting that the influence of the QBO is greatest in early winter. Using the same model, Gray (2003) found similar results, that the direction of the equatorial wind in the lower stratosphere influenced the early winter evolution but the equatorial upper stratosphere was more influential in mid- to late winter, in agreement with previous studies (e.g., Labitzke 1987; LvL88; Dunkerton and Baldwin 1992; Naito and Hirota 1997). There are two to three protracted migration events per year for the subtropical westerly jet. Poleward/downward migration by wave–mean flow interaction takes a fairly long time, consistent with the idea that the first event may be more affected by certain regions while the second or third, later in the winter, may be affected more by other regions and circumstances. There is also an increase in tropospheric wave forcing as the winter proceeds, which may be another influencing factor.

In the Smin/W composite (fourth column), the NH vortex is anomalously strong throughout the whole winter and an easterly anomaly does not appear until April, the time of the final warming. The Smax/W composite (second column) has a rather different pattern, however. Although the early winter is similarly quiet, in accordance with the Holton–Tan relationship since both are in westerly QBO phases, an easterly anomaly develops in February and moves poleward and downward by March, indicating midwinter warming events in Smax/W years despite the lack of a strong waveguide in the equatorial lower stratosphere. Thus, if one compares all of the December distributions, we see that the polar easterly anomalies are in Smin/E and Smax/E, as expected from the Holton–Tan relationship, but in February they are found in Smin/E and Smax/W. Hence a comparison of the two Smax composites in February shows a reversal of the Holton–Tan relationship. It is this feature of the SC– QBO interaction that has been most puzzling. However, we note that this feature is barely statistically significant at the 75% level. It may therefore be more accurate to say that the Holton–Tan relationship has been disrupted rather than reversed, at least until more data are available to improve statistical significance.

In Fig. 14 we show the daily time series of NP temperature at 50 hPa for each NH winter, again grouped into SC–QBO composites as described above. These plots not only show the day-to-day variations that are masked by the monthly averages but also provide information on the variability within each of the composites. The major features already noted in the monthly averaged plots of Fig. 13 are confirmed in Fig. 14 (see Table 2). In early winter (November–December), the most disturbed, and most variable, composites are the two QBO/E phase composites. In midwinter however (January–March), the two composites showing most disturbances in midwinter are the Smin/E and Smax/W composites. The disturbed nature of the Smax/W composite thus supports earlier evidence that the Holton– Tan relationship is disrupted in Smax years.

b. A speculative mechanism

We propose that the development of SSWs is sensitive to relatively small wind anomalies near the equatorial/ subtropical stratopause region, especially in mid- to late winter when the flow is highly nonlinear. We note that this altitude sensitivity near the tropical stratopause occurs where a meridional jet exists from inertial instability, as part of the semiannual oscillation (Hitchman and Leovy 1986), and the trajectory of planetary wave activity in latitude/altitude is such that the waves tend to focus on the tropical stratopause region [see Gray et al. (2003) and references therein]. Specifically, we propose that an easterly anomaly in this region helps to establish the Aleutian high, which then speeds up the onset of the SSW, while a westerly anomaly slows it down. In this way, the superposition of easterly anomalies associated with Smin conditions and QBO/E conditions results in SSWs in Smin/E years that are consistently earlier than average. Similarly, westerly anomalies associated with Smax conditions and QBO/W years superpose in Smax/W years and result in SSWs that are consistently later than average. In the other combinations that is, Smin/W and Smax/E the anomalies associated with the SC and QBO are of opposite sign and partially cancel, so there is no consistent speeding up or slowing down of the SSW in those years and hence no signal of SC–QBO interaction emerges.

Our results and those from earlier model experiments (Gray 2003) suggest that in early winter the influence of the lower-stratospheric equatorial winds proposed by Holton and Tan (1980, 1982) is strongest so that the polar vortex in early winter is more disturbed in QBO/E phase years. This is reflected in Table 2 with peak easterly anomalies in December of the QBO/E years, indicating a weakened vortex, and no corresponding signature in the QBO/W years. Later in the winter, however, when the flow is more highly nonlinear, the results suggest a greater influence from the equatorial and subtropical upper stratosphere, the region where both QBO and the 11-yr SC have zonal wind anomalies. In Smin/E both QBO and SC influences produce an easterly anomaly in that region and they will combine to produce a relatively strong easterly anomaly (see Table 1). The model results suggest that this will tend to speed up the onset of the SSWs. This is confirmed by the observations (Table 2) that indicate a consistently disturbed vortex in January. Similarly, in Smax/W years both QBO and SC influences produce a westerly anomaly in the upper equatorial stratosphere and these will reinforce each other and slow down the evolution of the SSWs. This is confirmed in Table 2, which indicates a consistently disturbed vortex in February, approximately 1 month later than in the Smin/E years. We note here that just because the onset of the SSWs is slowed down, this does not mean that the SSWs will not occur, provided there is time before the end of winter. Thus, in Smin/E and Smax/W years there will be a consistent pattern of early and late SSWs, respectively.

However, in Smin/W and Smax/E years the SC and QBO anomalies above 40 km are likely to partially cancel each other. Thus, there will be only a very weak influence on the SSW evolution from the upper equatorial stratosphere. In those years, with no strong guide from that region, the timing of the SSWs will not reflect any particular tendency toward either early or late SSWs; the SSWs will occur at any time throughout the winter and, when averaged together to form the composites, there will be no consistent pattern emerging, that is, no strong anomalies in any particular month. This again is confirmed in Table 2, since the main anomaly in both of those composites is in March/April at the time of the final warming.

Further research is required in order to fully test this proposed mechanism for SC–QBO influence on the NH polar vortex. One of the main problems is the short period of available data. This is crucial when compositing both on the basis of the 11-yr SC and the QBO. Statistical significance is therefore not high, and it may be many years before this is achieved. Nevertheless, we believe that a working hypothesis such as this may still be helpful. We also note the idealized nature of the model experiments, in terms of the model employed, the nature of the imposed forcing, and the lack of seasonality in the model runs. Although the recent GCM simulations of the SC–QBO interaction of Matthes et al. (2004) are encouraging and suggest that this mechanism is active, further GCM experiments would be useful.

Acknowledgments

We thank the European Centre for Medium-Range Weather Forecasts for the development of such a high-quality dataset and Ag Stevens of the British Atmospheric Data Centre for his help in acquiring it.

REFERENCES

  • Baldwin, M. P., and Coauthors, 2001: The quasi-biennial oscillation. Rev. Geophys, 39 , 179229.

  • Crooks, S., , and L. J. Gray, 2004: An examination of the influence of the 11-year solar cycle on stratospheric climate using the ERA-40 data. J. Climate, submitted.

    • Search Google Scholar
    • Export Citation
  • Dunkerton, T. J., , and M. P. Baldwin, 1992: Modes of interannual variability in the stratosphere. Geophys. Res. Lett, 19 , 4952.

  • Gray, L. J., 2003: The influence of the equatorial upper stratosphere on stratospheric sudden warmings. Geophys. Res. Lett.,30, 1166, doi:10.1029/2002GL016430.

    • Search Google Scholar
    • Export Citation
  • Gray, L. J., , E. F. Drysdale, , T. J. Dunkerton, , and B. N. Lawrence, 2001a: Model studies of the interannual variability of the northern hemisphere stratospheric winter circulation: The role of the quasi-biennial oscillation. Quart. J. Roy. Meteor. Soc, 127 , 14131432.

    • Search Google Scholar
    • Export Citation
  • Gray, L. J., , S. J. Phipps, , T. J. Dunkerton, , M. P. Baldwin, , E. F. Drysdale, , and M. R. Allen, 2001b: A data study of the influence of the equatorial upper stratosphere on northern hemisphere stratospheric sudden warmings. Quart. J. Roy. Meteor. Soc, 127 , 19852003.

    • Search Google Scholar
    • Export Citation
  • Gray, L. J., , S. Sparrow, , M. Juckes, , A. O'Neill, , and D. G. Andrews, 2003: Flow regimes in the winter stratosphere of the northern hemisphere. Quart. J. Roy. Meteor. Soc, 129 , 925945.

    • Search Google Scholar
    • Export Citation
  • Haigh, J. D., 2003: The effect of solar variability on the Earth's climate. Philos. Trans. Roy. Soc. London, 361 , 95111.

  • Harvey, V. L., , M. H. Hitchman, , R. B. Pearce, , and T. D. Fairly, 1999: Tropical aerosol in the Aleutian High. J. Geophys. Res, 104 (D6) 62816290.

    • Search Google Scholar
    • Export Citation
  • Haynes, P. H., 1985: Nonlinear instability of a Rossby-wave critical layer. J. Fluid Mech, 161 , 493511.

  • Hitchman, M. H., , and C. B. Leovy, 1986: Evolution of the zonal mean state in the equatorial middle stratosphere during October 1978–May 1979. J. Atmos. Sci, 43 , 31593176.

    • Search Google Scholar
    • Export Citation
  • Holton, J. R., , and H-C. Tan, 1980: The influence of the equatorial quasi-biennial oscillation on the global circulation at 50 mb. J. Atmos. Sci, 37 , 22002208.

    • Search Google Scholar
    • Export Citation
  • Holton, J. R., , and H-C. Tan, 1982: The quasi-biennial oscillation in the Northern Hemisphere lower stratosphere. J. Meteor. Soc. Japan, 60 , 140148.

    • Search Google Scholar
    • Export Citation
  • Hood, L. L., 2004: Effects of solar UV variability on the stratosphere. Solar Variability and Its Effect on the Earth's Atmosphere and Climate System. Geophys. Monogr, No. 141, Amer. Geophys. Union,. 283303.

    • Search Google Scholar
    • Export Citation
  • Hood, L. L., , and B. E. Soukharev, 2003: Quasi-decadal variability of the tropical lower stratosphere: The role of extratropical wave forcing. J. Atmos. Sci, 60 , 23892403.

    • Search Google Scholar
    • Export Citation
  • Killworth, P. D., , and M. E. McIntyre, 1985: Do Rossby-wave critical layers absorb, reflect or overreflect? J. Fluid Mech, 161 , 449492.

    • Search Google Scholar
    • Export Citation
  • Kodera, K., , and Y. Kuroda, 2002: Dynamical response to the solar cycle: Winter stratopause and lower stratosphere. J. Geophys. Res.,107, 4749, doi:10.1029/2002JD002224.

    • Search Google Scholar
    • Export Citation
  • Kodera, K., , Y. Kuroda, , and S. Pawson, 2000: Stratospheric sudden warmings and slowly propagating zonal-mean zonal wind anomalies. J. Geophys. Res, 105 , 1235112359.

    • Search Google Scholar
    • Export Citation
  • Labitzke, K., 1987: Sunspots, the QBO and the stratospheric temperature in the north polar region. Geophys. Res. Lett, 14 , 535537.

  • Labitzke, K., 2003: The global signal of the 11-year sunspot cycle in the atmosphere: When do we need the QBO? Meteor. Z, 12 , 209216.

    • Search Google Scholar
    • Export Citation
  • Labitzke, K., , and H. van Loon, 1988: Association between the 11-year solar cycle, the QBO and the atmosphere. Part I: The troposphere and stratosphere in northern hemisphere in winter. J. Atmos. Terr. Phys, 50 , 197206.

    • Search Google Scholar
    • Export Citation
  • Labitzke, K., , and H. van Loon, 1999: The Stratosphere: Phenomena, History and Relevance. Springer-Verlag, 179 pp.

  • Labitzke, K., , and H. van Loon, 2000: The QBO effect on the global stratosphere in northern winter. J. Atmos. Sol.-Terr. Phys, 62 , 621628.

    • Search Google Scholar
    • Export Citation
  • Lindzen, R. S., , and K. K. Tung, 1978: Wave overreflection and shear instability. J. Atmos. Sci, 35 , 16261632.

  • Matthes, K., , U. Langematz, , L. J. Gray, , K. Kodera, , and K. Labitzke, 2004: Improved 11-year solar signal in the FUB-CMAM. J. Geophys. Res.,109, D06101, doi:10.1029/2003JD004012.

    • Search Google Scholar
    • Export Citation
  • McCormack, J. P., 2003: The influence of the 11-year solar cycle on the quasi-biennial oscillation. Geophys. Res. Lett.,30, 2162, doi:10.1029/2003GL018314.

    • Search Google Scholar
    • Export Citation
  • Naito, Y., , and I. Hirota, 1997: Interannual variability of the northern winter stratospheric circulation related to the QBO and the solar cycle. J. Meteor. Soc. Japan, 75 , 925937.

    • Search Google Scholar
    • Export Citation
  • O'Neill, A., , and V. Pope, 1988: Simulations of linear and nonlinear disturbances in the stratosphere. Quart. J. Roy. Meteor. Soc, 114 , 10631110.

    • Search Google Scholar
    • Export Citation
  • O'Sullivan, D., , and T. J. Dunkerton, 1994: Seasonal development of the extratropical QBO in a numerical model of the middle atmosphere. J. Atmos. Sci, 51 , 37063721.

    • Search Google Scholar
    • Export Citation
  • Polvani, L. M., , and R. Saravanan, 2000: The three dimensional structure of breaking Rossby waves in the polar wintertime stratosphere. J. Atmos. Sci, 57 , 36633685.

    • Search Google Scholar
    • Export Citation
  • Robinson, W. A., 1986: The behavior of planetary wave 2 in preconditioned zonal flows. J. Atmos. Sci, 43 , 31093121.

  • Salby, M., , and P. Callaghan, 2000: Connection between the solar cycle and the QBO: The missing link. J. Climate, 13 , 328338.

  • Scaife, A. A., , and I. N. James, 2000: Response of the stratosphere to interannual variability of tropospheric planetary waves. Quart. J. Roy. Meteor. Soc, 126 , 275297.

    • Search Google Scholar
    • Export Citation

Fig. 1.
Fig. 1.

(a) Time series of zonally averaged, J–F-averaged NP temperatures (K) at 30 hPa (approximately 24 km) (solid line) and zonally averaged zonal wind (m s−1) at the equator at 44 hPa (dotted line). (b) Time series of 10.7-cm solar radio flux (10−22 W m−2 Hz−1). (c) Time series of zonally averaged, J–F-averaged NP temperatures (K) at 30 hPa (solid line) for years in which the QBO was in its west phase. The 10.7-cm solar flux (10−22 W m−2 Hz−1) is also superposed (dotted lines). (d) As in (c) but for the QBO/E phase years

Citation: Journal of the Atmospheric Sciences 61, 23; 10.1175/JAS-3297.1

Fig. 2.
Fig. 2.

QBO composites: monthly averaged latitude–height sections of zonally averaged zonal wind anomalies (m s−1) for the period 1979–2001 in the months Oct–Apr. The data have been composited into (left) QBO/E and (middle) QBO/W winters, as described in the text. The right-hand column shows the QBO/E minus QBO/W difference. Contour interval is 2 m s−1, dashed contour is zero, and dotted contours are negative values. In the left and middle columns, shaded regions indicate values less than −4 m s−1. In the right-hand column, the light and dark shading indicate statistical confidence at the 75% and 95% levels, respectively, using a standard Student's t test

Citation: Journal of the Atmospheric Sciences 61, 23; 10.1175/JAS-3297.1

Fig. 3.
Fig. 3.

As in Fig. 2 but for SC composites. The data have been composited into (left) Smin and (middle) Smax years, as described in the text. The right-hand column shows the Smax minus Smin difference

Citation: Journal of the Atmospheric Sciences 61, 23; 10.1175/JAS-3297.1

Fig. 4.
Fig. 4.

Time series of area-weighted, zonally averaged modeled temperatures (K) north of 62.5°N at 32 km from (a) the control run, (b) E20–30, (c) E25–35, (d) E30–40, (e) E35–45, and (f) E40–50

Citation: Journal of the Atmospheric Sciences 61, 23; 10.1175/JAS-3297.1

Fig. 5.
Fig. 5.

Ensemble-averaged Eliassen–Palm flux diagnostics for experiments (left) E25–35 and (middle) E30–40 and (right) their difference (E25– 35 minus E30–40) at various days throughout the integration. The contours of E–P flux divergence are in m s−1, according to the “acceleration” scaling described in Gray et al. (2003). Solid contours are positive, the dashed contour is zero, and dotted contours are negative. The shading in the first two columns highlights regions less than −2 m s−1. The shading in the difference plots indicates significance at the 99% level

Citation: Journal of the Atmospheric Sciences 61, 23; 10.1175/JAS-3297.1

Fig. 6.
Fig. 6.

Time series of area-weighted, zonally averaged temperature (K) north of 62.5°N at various heights from a typical ensemble member of the (a) E25–35 experiment and (b) E30–40 experiment

Citation: Journal of the Atmospheric Sciences 61, 23; 10.1175/JAS-3297.1

Fig. 7.
Fig. 7.

Time series of area-weighted, zonally averaged temperature (K) at 44 km at each latitude grid box from (a) E25–35 and (b) E30– 40. The plots span from 2.5° (top line) to 87.5°N (bottom line) in 5° latitude intervals

Citation: Journal of the Atmospheric Sciences 61, 23; 10.1175/JAS-3297.1

Fig. 8.
Fig. 8.

Typical latitude–height sections of zonally averaged zonal winds (m s−1) on day 10 from (top) S40–50 experiment, (middle) the control experiment, and (bottom) the difference between them

Citation: Journal of the Atmospheric Sciences 61, 23; 10.1175/JAS-3297.1

Fig. 9.
Fig. 9.

As in Fig. 4 but from (a) the S40–50 ensemble and (b) the control ensemble

Citation: Journal of the Atmospheric Sciences 61, 23; 10.1175/JAS-3297.1

Fig. 10.
Fig. 10.

Typical distribution of NH Ertel potential vorticity on the 1650-K isentropic surface (10−4 K m−1 s−1) at day 50 from (a) the S40–50 experiment and (b) the control experiment

Citation: Journal of the Atmospheric Sciences 61, 23; 10.1175/JAS-3297.1

Fig. 11.
Fig. 11.

As in Fig. 5 but for (left) experiment S40–50 and (middle) the control experiment and (right) their difference (S40–50 minus control) at various days throughout the integration. The arrows in the top two rows have been magnified by a factor of 20 and 30, respectively, compared with the remaining plots to improve legibility

Citation: Journal of the Atmospheric Sciences 61, 23; 10.1175/JAS-3297.1

Fig. 12.
Fig. 12.

Typical latitude–height sections of scaled Ertel potential vorticity (as described in Polvani and Saravanan 2000) at day 20 from (left) the S40–50 experiment, (middle) the control experiment, and (right) the difference between them (10−4 K m−1 s−1)

Citation: Journal of the Atmospheric Sciences 61, 23; 10.1175/JAS-3297.1

Fig. 13.
Fig. 13.

Monthly averaged latitude–height sections of zonally averaged zonal wind anomalies (m s−1) for the period 1958–2001 in the months Oct–Apr. The data have been composited into (first column) Smin/E, (second column) Smax/W, (third column) Smax/E, and (fourth column) Smin/W years, as described in the text. Contour interval is 5 m s−1 with additional ±2 m s−1 contours also included. Solid contours are positive; dashed contour is zero. Dotted contours are negative values. The light and dark shading indicate statistical confidence at the 75% and 95% levels, respectively. These estimates are for data values different from zero, using a standard Student's t test

Citation: Journal of the Atmospheric Sciences 61, 23; 10.1175/JAS-3297.1

Fig. 14.
Fig. 14.

Time series of daily averaged NP temperature (K) at 50 hPa. The data have been composited into Smin/E, Smin/W, Smax/E and Smax/W years, as described in the text

Citation: Journal of the Atmospheric Sciences 61, 23; 10.1175/JAS-3297.1

Table 1.

Summary of the phase of the zonal wind anomaly in the upper equatorial stratosphere (>40 km) associated with the QBO, the 11-yr SC, and the nature of the anomaly when they are added together. Also indicated in bold are the predicted effects of this anomaly on the timing of sudden stratospheric warmings, based on the results of the model experiments

Table 1.
Table 2.

Summary of the months in which the ERA-40 data show significant easterly anomalies in Fig. 13, indicating the presence of warming events

Table 2.
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