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  • View in gallery

    Distribution of total ozone (DU) during January: (a) computed by 3D model and (b) climatological mean observed by TOMS.

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    Anomalous wintertime tendency of temperature (K per season) during early winter (October–January), forced by a one-standard-deviation intensification of EP flux from the troposphere.

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    Anomalous residual mean circulation under conditions of Fig. 2.

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    As in Fig. 2, but for the wintertime tendency of zonal wind (m s−1 per season).

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    Montgomery streamfunction during January at θ = 1018 K for (a) a one-standard-deviation intensification of EP flux from the troposphere and (b) a one-standard-deviation weakening of EP flux from the troposphere.

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    Anomalous wintertime tendency of temperature (K per season) during late winter (November–March), forced by a one-standard-deviation intensification of EP flux from the troposphere and, simultaneously, an easterly swing of the QBO.

  • View in gallery

    As in Fig. 6, but for the wintertime tendency of ozone mixing ratio (ppmv per season).

  • View in gallery

    As in Fig. 6, but for the wintertime tendency of total ozone (DU per season) in the presence of homogeneous/gas-phase chemistry (solid) and inclusive of heterogeneous processes (dashed).

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    As in Fig. 6, but for the QBO contribution alone (EP flux fixed).

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    As in Fig. 9, but for the wintertime tendency of total ozone.

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    Anomalous seasonal-mean tendency of column-averaged temperature above θ = 360 K, as function of central month, averaged poleward of 60°N (solid) and from 60°N to 30°S (dashed). Column average corresponds to vertical integral over pressure, making it analogous to total ozone. Seasonal-mean tendency calculated as sliding 3-month difference.

  • View in gallery

    As in Fig. 11, but for the seasonal-mean tendency of total ozone.

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Interannual Changes of Stratospheric Temperature and Ozone: Forcing by Anomalous Wave Driving and the QBO

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  • 1 Macquarie University, Sydney, New South Wales, Australia
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Abstract

A 3D model of dynamics and photochemistry is used to investigate interannual changes of stratospheric dynamical and chemical structure through their dependence on tropospheric planetary waves and on the quasi-biennial oscillation (QBO). The integrations reproduce the salient features of the climate sensitivities of temperature and ozone, which have been composited from the observed records of ECMWF and the Total Ozone Mapping Spectrometer (TOMS). Characterized by a strong anomaly of one sign at polar latitudes and a comparatively weak anomaly of opposite sign at subpolar latitudes, each bears the signature of the residual mean circulation. The structure is very similar to that associated with the Arctic Oscillation.

The integrations imply that, jointly, anomalous Eliassen–Palm (EP) flux transmitted from the troposphere by planetary waves and the QBO are the major mechanisms behind interannual changes in the stratosphere. An analogous conclusion follows from the observational record. During early winter, anomalous temperature and ozone are accounted for almost entirely by anomalous EP flux from the troposphere, as they are in the observational record. During late winter, both mechanisms are required to reproduce observed anomalies. Although the QBO forces anomalous structure equatorward of 40°N, the strong anomaly over the Arctic follows principally from anomalous upward EP flux. Reflecting anomalous wave driving of residual mean motion, the change of EP flux leads to anomalous downwelling of ozone-rich air. In concert with isentropic mixing by planetary waves, the anomalous enrichment that ensues at extratropical latitudes sharply modifies total ozone over the Arctic. Integrations distinguished by the omission of heterogeneous processes indicate that chemical destruction accounts for approximately 20% of the anomaly in Arctic ozone between warm and cold winters. Analogous to estimates derived from the observed record of the Solar Backscatter Ultraviolet, version 8 (SBUV-V8) instrument, the remaining approximately 80% follows from anomalous transport.

The climate sensitivities of temperature and ozone describe random changes between years, introduced by anomalous EP flux and the QBO. Those interannual changes evolve with a particular seasonality. Like their structure, the seasonal dependence of anomalous temperature and ozone bears the signature of the residual mean circulation. Systematic changes in the observed record, which comprise stratospheric trends, have similar structure and seasonality.

Corresponding author address: Murry L. Salby, Environmental Science, Macquarie University, Sydney NSW 2109, Australia. E-mail: pcd@mq.edu.au

Abstract

A 3D model of dynamics and photochemistry is used to investigate interannual changes of stratospheric dynamical and chemical structure through their dependence on tropospheric planetary waves and on the quasi-biennial oscillation (QBO). The integrations reproduce the salient features of the climate sensitivities of temperature and ozone, which have been composited from the observed records of ECMWF and the Total Ozone Mapping Spectrometer (TOMS). Characterized by a strong anomaly of one sign at polar latitudes and a comparatively weak anomaly of opposite sign at subpolar latitudes, each bears the signature of the residual mean circulation. The structure is very similar to that associated with the Arctic Oscillation.

The integrations imply that, jointly, anomalous Eliassen–Palm (EP) flux transmitted from the troposphere by planetary waves and the QBO are the major mechanisms behind interannual changes in the stratosphere. An analogous conclusion follows from the observational record. During early winter, anomalous temperature and ozone are accounted for almost entirely by anomalous EP flux from the troposphere, as they are in the observational record. During late winter, both mechanisms are required to reproduce observed anomalies. Although the QBO forces anomalous structure equatorward of 40°N, the strong anomaly over the Arctic follows principally from anomalous upward EP flux. Reflecting anomalous wave driving of residual mean motion, the change of EP flux leads to anomalous downwelling of ozone-rich air. In concert with isentropic mixing by planetary waves, the anomalous enrichment that ensues at extratropical latitudes sharply modifies total ozone over the Arctic. Integrations distinguished by the omission of heterogeneous processes indicate that chemical destruction accounts for approximately 20% of the anomaly in Arctic ozone between warm and cold winters. Analogous to estimates derived from the observed record of the Solar Backscatter Ultraviolet, version 8 (SBUV-V8) instrument, the remaining approximately 80% follows from anomalous transport.

The climate sensitivities of temperature and ozone describe random changes between years, introduced by anomalous EP flux and the QBO. Those interannual changes evolve with a particular seasonality. Like their structure, the seasonal dependence of anomalous temperature and ozone bears the signature of the residual mean circulation. Systematic changes in the observed record, which comprise stratospheric trends, have similar structure and seasonality.

Corresponding author address: Murry L. Salby, Environmental Science, Macquarie University, Sydney NSW 2109, Australia. E-mail: pcd@mq.edu.au

1. Introduction

The wintertime circulation of the Northern Hemisphere stratosphere changes significantly between years. Accompanying changes of dynamical structure are changes of Northern Hemisphere ozone. Inherent to each is the residual mean circulation, which shapes thermal and chemical structure. Through adiabatic warming and poleward transport, the Brewer–Dobson circulation controls how cold the temperature becomes during an individual winter, as well as how much total ozone increases (Brewer 1949; Dobson 1956; Murgatroyd and Singleton 1961).

The residual circulation is driven by planetary waves that transmit momentum upward from the troposphere. When absorbed, that momentum drives a poleward drift, which converges at high latitude to force mean downwelling. The accompanying adiabatic warming maintains Arctic temperature as warmer and the polar-night vortex as weaker than each would be under conditions of radiative equilibrium. Compensating that vertical motion at lower latitudes is mean upwelling, which is accompanied by adiabatic cooling. Simultaneously, the poleward drift transfers ozone-rich air from its chemical source at low latitude into the winter hemisphere, where total ozone increases during winter by as much as 100%.

These consequences of the residual circulation are closely related to planetary wave activity, which transmits momentum upward from the troposphere. Measured by the net upward Eliassen–Palm (EP) flux near the tropopause , the momentum delivered to the middle atmosphere during winter is determined chiefly by planetary wave structure in the troposphere. An intensification of tropospheric planetary waves leads to anomalous upward EP flux, which, upon being absorbed in the middle atmosphere, forces intensified residual mean motion. In concert with intensified isentropic mixing by planetary waves, those conditions favor a polar-night vortex that is anomalously warm and weak. Conversely, a weakening of tropospheric planetary waves favors a polar-night vortex that is anomalously cold and strong.

A similar influence is exerted by the quasi-biennial oscillation (QBO) of stratospheric equatorial wind uEQ. By displacing the critical line of planetary waves, the QBO controls where in the middle atmosphere EP flux is absorbed and, hence, where residual mean motion is forced (Holton and Tan 1980; McIntyre 1982). During QBO easterlies, the critical line and mean meridional motion are advanced into the winter hemisphere, along with downwelling and adiabatic warming. Those conditions favor a polar-night vortex that is anomalously warm and weak, with increased wintertime ozone (Gray and Pyle 1989; O’Sullivan and Salby 1990; O’Sullivan and Young 1992; Tung and Yang 1994). During QBO westerlies, the critical line and mean meridional motion are removed into the summer hemisphere. Those conditions favor a polar-night vortex that is anomalously cold and strong, with reduced wintertime ozone.

Jointly, changes of upward EP flux from the troposphere and of equatorial wind associated with the QBO represent anomalous forcing of the residual circulation. They drive an anomalous residual circulation, one that modulates climatological-mean residual motion. Implied are commensurate changes of wintertime temperature and ozone. Augmenting these dynamical influences are sporadic influences from volcanic eruptions and ENSO, as well as the solar cycle. For interannual changes collected over several decades, however, the latter (if only because they are infrequent or cyclical) are secondary to the influences of EP flux and the QBO. Indeed, observed forcing of the residual circulation bears a close relationship to observed changes of the polar-night vortex, as well as to changes of Northern Hemisphere ozone (Fusco and Salby 1999). Each tracks observed changes of upward EP flux from the troposphere. Supported by changes of equatorial wind associated with the QBO, changes of EP flux account for much of the interannual variance of wintertime temperature and ozone (Hadjinicolaou et al. 1997, 2002, 2005; Salby and Callaghan 2002; Hu and Tung 2002).

Interannual changes of temperature and ozone are chiefly random. Nevertheless, they will be seen shortly to have a concrete relationship to major dynamical influences, with which they operate coherently. The climate sensitivities of temperature and ozone to changes in those controlling influences have been determined from a large population of winters (Salby and Callaghan 2002). They have structure that is coherent but out of phase between high and low latitudes. The compensating changes reflect anomalous downwelling and upwelling, which are inherent to the residual circulation. Also reflecting the residual circulation is the seasonality of anomalous temperature and ozone. Both amplify during the disturbed season, when the residual circulation is strong. They then decay following the final warming, when the residual circulation collapses.

Characterizing random changes between years, the foregoing structure and seasonality operate coherently with anomalous forcing of the residual circulation. Having much the same structure and seasonality are systematic changes, which comprise trends of stratospheric temperature and ozone (Salby and Callaghan 2004a).

In addition to anomalous transport, changes of ozone also derive from anomalous photochemistry, which is coupled to the residual circulation through temperature. The temperature dependence of heterogeneous processes makes anomalous photochemistry important at polar latitudes. There, chlorine activation through polar stratospheric clouds (PSCs) and aerosol can amplify ozone changes that are introduced through anomalous transport. In fact, like other features of the wintertime circulation, PSC and chlorine activation vary coherently with anomalous forcing of the residual circulation (Weber et al. 2003). Anomalously weak forcing of residual mean motion leads to anomalously cold conditions, which are attended by increased PSC and chlorine activation. Accompanying reduced ozone transport, the latter also imply reduced ozone through accelerated chemical destruction. The chemical contribution to anomalous Northern Hemisphere ozone is thus of the same sense as the dynamical contribution. The magnitude of this contribution remains uncertain. However, it appears to be nonessential at midlatitudes, where observed changes are largely reproduced by integrations with a chemical transport model (CTM) exclusive of heterogeneous processes (Hadjinicolaou et al. 2002).

To elucidate the mechanisms behind observed changes, we employ a 3D model of dynamics and photochemistry. In it, we explore the dependence of stratospheric dynamical and chemical structure on tropospheric planetary waves and on the QBO. Following an overview of the model, section 3 calculates the climate sensitivity of Northern Hemisphere temperature and ozone to changes of tropospheric wave structure and of equatorial wind. Reflecting the response to random changes between years, the computed climate sensitivities are shown to reproduce the salient structure of those that have been composited from the observational record. Anomalous ozone, which accounts for homogeneous chemistry as well as temperature-dependent heterogeneous processes, is shown to follow principally from anomalous transport. This conclusion applies even over the Arctic, where anomalous transport is augmented by anomalous chemical destruction associated with PSC and aerosol. Section 4 then evaluates the seasonality of anomalous temperature and ozone. Like anomalous structure, the computed seasonality reproduces the major features of seasonality that have been composited from the observational record. These characteristics of random interannual changes are then related in section 5 to systematic changes that comprise trends of temperature and ozone.

2. Numerical framework

a. Model of dynamics and photochemistry

The 3D model integrates the nonlinear primitive equations through spectral technique, along with a family treatment of photochemistry (Callaghan et al. 1999; Fusco and Salby 1999). Formulated in isentropic coordinates, the model extends from an isentropic surface in the upper troposphere, corresponding to θ = 360 K (~200 mb), through the mesosphere. It is capped overhead by a deep thermal sponge layer that achieves the radiation condition by absorbing wave activity above 85 km.

Coupled to the lower boundary is a reservoir layer, which serves as an amorphous lower troposphere by enforcing conservation of mass. Air rejected from the stratosphere at high latitude is compensated at lower latitude by air that is absorbed from the troposphere at the same rate. Inside the reservoir layer is collected, among other species, stratospheric ozone that has been transferred across the lower boundary through mean downwelling. After entering the lower troposphere, stratospheric ozone is destroyed on the time scale of a season, representative of oxidation processes and the observed seasonality of tropospheric ozone (London 1985).

The family treatment of photochemistry calculates the distributions of some 50 species that influence ozone (Fusco 1997; Fusco and Salby 1999). It has been reformulated through an asymptotic partitioning of chemical species, wherein the controlling reactions are determined recursively from the fastest reaction rates. This formal treatment achieves a robust description of chemical species that have widely varying lifetimes. While retaining the computational advantage of the family representation, the asymptotic partitioning of species remains well behaved in 3D integrations, wherein diurnally varying photolysis enables short-lived species to undergo a strong diurnal variation.

Applied in tandem with the family treatment of photochemistry is an explicit treatment of individual reactions by version 3 of the Model for Ozone and Related Tracers (MOZART-3) (Brasseur et al. 1998; Horowitz et al. 2003). Included is a comprehensive representation of stratospheric aerosol. Calculated explicitly are the surface area densities of sulfate aerosol, nitric acid trihydrate (NAT), and PSC ice. Like the family treatment, MOZART-3 is married with the dynamical integration. Jointly, those calculations are used to isolate the contribution to anomalous ozone from heterogeneous processes.

Forced at its lower boundary by observed tropospheric behavior, the 3D model reproduces major features of dynamical and chemical structure in the wintertime stratosphere. Included are zonal-mean motion, the residual circulation, and ozone, as well as changes associated with each (Francis and Salby 2001; Callaghan and Salby 2002; Salby et al. 2002).

b. Climate sensitivity analysis

In the present study, parallel integrations are performed to evaluate the dependence of dynamical and chemical structure on anomalous forcing of the residual circulation. Upward EP flux from the troposphere undergoes an annual cycle, maximizing around solstice. However, daily values of vary between years, comprising a population of annual cycles (Salby and Callaghan 2002). To explore the general character of interannual differences and how they influence dynamical and chemical structure, the population of annual cycles is considered in two groups. One has a mean equal to the climatological-mean annual cycle of plus one standard deviation; the other has a mean equal to the climatological-mean annual cycle of minus one standard deviation. Consolidating annual cycles within each group leads to a cancellation of individual wave events, which operate on a time scale of days and are incoherent between years. This eliminates day-to-day differences of between the groups, in favor of the coherent interannual anomaly between them.1 Separated by two standard deviations, the group-mean annual cycles of are mutually distinct and support direct comparison with observed diagnostics. Parallel integrations forced by the two perturbed annual cycles provide a direct counterpart of anomalous structure that has been composited from the observed record (Salby and Callaghan 2002), against which the integrations can then be compared.

An integration in which tropospheric planetary waves are amplified yields structure under conditions of intensified . An integration in which tropospheric planetary waves are suppressed yields structure under conditions of weakened . Differencing the integrations and dividing by 2 yields the anomalous dynamical and chemical structures that are introduced by a one-standard-deviation intensification of , which define the respective climate sensitivities.

Similar integrations are performed for extremal phases of the QBO, which modulates residual motion through interaction with planetary waves. The QBO is imposed through a momentum source that drives equatorial wind toward the average structure of the QBO for a specified phase, as has been composited from observations; see Garcia and Salby (1987) for a detailed description. Represented analytically, the imposed QBO is concentrated within 20° of the equator, intensifies upward above the tropopause, and then decays gradually above 40 km, where the semiannual oscillation (SAO) prevails. Parallel integrations are again performed, one with the QBO in its easterly phase and another with the QBO in its westerly phase. Differencing those integrations and dividing by 2 then obtains the anomalous dynamical and chemical structures that are introduced by the swing of uEQ, which defines the respective climate sensitivities.2

3. Anomalous temperature and ozone

Forming the control is a seasonal integration, from autumnal equinox through spring. In it, tropospheric wave activity is prescribed at the lower boundary from the record of European Centre for Medium-Range Weather Forecasts (ECMWF) reanalyses: averaging individual calendar days over 1970–2002 in the 40-yr ECMWF Re-Analysis (ERA-40) record gives the climatological-mean seasonal variation. Integrations forced by this imposed tropospheric wave activity are ramped up from undisturbed conditions during summer. They reproduce the major features of observed dynamical structure in the wintertime stratosphere (see, e.g., Callaghan et al. 1999; Francis and Salby 2001; Callaghan and Salby 2002).

Figure 1 compares the computed distribution of total ozone during January in the control against the climatological-mean distribution of 〈O3〉 observed by the Total Ozone Mapping Spectrometer (TOMS) over 1979–2002. In each, 〈O3〉 is marked by crescent structure, wherein high ozone circumscribes the Arctic. In the control (Fig. 1a), total ozone maximizes over the Orient, where 〈O3〉 approaches 500 Dobson units (DU). Much the same structure appears in the observed distribution (Fig. 1b). Along with secondary maxima over Europe and North America and a minimum over the Barents Sea, the computed distribution of total ozone reproduces the major features of the climatological-mean distribution observed by TOMS.

Fig. 1.
Fig. 1.

Distribution of total ozone (DU) during January: (a) computed by 3D model and (b) climatological mean observed by TOMS.

Citation: Journal of the Atmospheric Sciences 68, 7; 10.1175/2011JAS3671.1

a. Early winter

Considered first are the months October–January, the interval during early winter that determines the wintertime minimum of Northern Hemisphere temperature. We then examine changes in the wintertime decrease of temperature
e1
where ΔTJan–Oct represents the wintertime-mean tendency of temperature. In the observational record, its interannual variance during this season is largely accounted for by anomalous alone (Salby and Callaghan 2002). The net upward EP flux, in turn, is controlled principally by wave structure in the troposphere.3

Interannual variance of increases during winter, from a standard deviation of about 15% during autumn to about 30% near solstice (Salby and Callaghan 2002). Two parallel integrations are performed. In one, the climatological-mean seasonal variation of tropospheric wave structure is amplified uniformly, intensifying by 20%. In the other, it is suppressed uniformly, weakening by 20%. Differencing the integrations and dividing by 2 then recovers the corresponding climate sensitivity of temperature: the anomalous temperature introduced by a one-standard-deviation intensification of .

Intensified introduces anomalous downwelling and adiabatic warming. Through the thermodynamic equation, the latter modifies the wintertime tendency of temperature. Plotted in Fig. 2, as a function of latitude and altitude, is the anomalous wintertime tendency of temperature during early winter [Eq. (1)]. Recovered from the parallel integrations, ΔTJan–Oct is the computational analog of the anomalous wintertime tendency that was composited from the ECMWF record in Salby and Callaghan (2002). The anomalous temperature tendency is marked by a strong warm anomaly over the Arctic. There, ΔTJan–Oct increases upward, exceeding 7 K per season near 35 km (~5 mb). Equatorward of 50°N, anomalous ΔTJan–Oct reverses. The cold anomaly extends across the tropics and into the summer hemisphere. However, it is an order of magnitude weaker that the warm anomaly over the Arctic. Much the same structure characterizes the observed climate sensitivity of temperature during early winter, which likewise approaches 7 K per season (Salby and Callaghan 2002; see their Fig. 5).

Fig. 2.
Fig. 2.

Anomalous wintertime tendency of temperature (K per season) during early winter (October–January), forced by a one-standard-deviation intensification of EP flux from the troposphere.

Citation: Journal of the Atmospheric Sciences 68, 7; 10.1175/2011JAS3671.1

In the mesosphere, which lies above the conventional observing network, the anomalous temperature tendency reverses. Capping the warm anomaly in the Arctic stratosphere is a cold anomaly. Likewise, the cold anomaly in the subpolar stratosphere is capped overhead by a warm anomaly, one that occupies the same range of latitude.

These features of anomalous temperature derive from anomalous residual mean motion. Intensified introduces anomalous · F and zonal-mean drag. The Coriolis torque then drives anomalous poleward motion, which in turn leads to anomalous downwelling over the winter pole. Figure 3 plots anomalous residual mean motion recovered from the parallel integrations. Clearly visible in the wintertime upper stratosphere is anomalous poleward motion . Upon converging over the Arctic, it drives anomalous downwelling . The attending adiabatic warming leads to the strong positive anomaly in the wintertime tendency of temperature (Fig. 2). Compensating anomalous downwelling is anomalous upwelling over the tropics and summer hemisphere. Visibly weaker, it accounts for the negative, but comparatively weak, temperature anomaly found at the same latitudes.

Fig. 3.
Fig. 3.

Anomalous residual mean circulation under conditions of Fig. 2.

Citation: Journal of the Atmospheric Sciences 68, 7; 10.1175/2011JAS3671.1

In the mesosphere is an anomalous circulation of opposite sense: There, anomalous poleward motion near the stratopause converges over the Arctic to drive anomalous upwelling. Compensating it over the tropics and summer hemisphere is anomalous downwelling. As in the stratosphere, adiabatic cooling and warming that attend these vertical motions account for the anomalous temperature tendencies that are found in those regions.

Plotted in Fig. 4 is the anomalous wintertime tendency of zonal wind derived from the parallel integrations. Note that ΔuJan–Oct is strong and negative near 60°N, where it reflects a deceleration of the polar-night jet during winters of intensified . The computed structure of ΔuJan–Oct approaches −20 m s−1 per season near the stratopause. Reversed values appear at lower latitude. The climate sensitivity of zonal wind mirrors that which was composited from observed winters in the ECMWF record (Salby and Callaghan 2002; see their Fig. 6).

Fig. 4.
Fig. 4.

As in Fig. 2, but for the wintertime tendency of zonal wind (m s−1 per season).

Citation: Journal of the Atmospheric Sciences 68, 7; 10.1175/2011JAS3671.1

Compared in Fig. 5 is the January Montgomery streamfunction at θ = 1018 K (near 10 mb) in the presence of amplified and suppressed tropospheric wave activity. Under conditions of intensified (Fig. 5a), the Aleutian high is centered near 55°N. Anticyclonic motion just invades the Arctic. In the opposite hemisphere, the polar-night vortex has been displaced equatorward along the Greenwich meridian, centered near 80°N. Under conditions of weakened (Fig. 5b), the Aleutian high is weaker and centered some 10° equatorward, near 45°N. The cyclonic flow about the Arctic is therefore left relatively undisturbed, with the polar-night vortex broader, deeper, and nearly centered over the pole.

Fig. 5.
Fig. 5.

Montgomery streamfunction during January at θ = 1018 K for (a) a one-standard-deviation intensification of EP flux from the troposphere and (b) a one-standard-deviation weakening of EP flux from the troposphere.

Citation: Journal of the Atmospheric Sciences 68, 7; 10.1175/2011JAS3671.1

b. Late winter

Considered next are the months November–March, the interval during late winter that determines the spring maximum of total ozone. As in the observed record, we consider changes in the wintertime increase of total ozone. In the TOMS record, interannual variance of ozone is still accounted for principally by anomalous . However, during late winter, anomalous receives support from anomalous uEQ associated with the QBO (Salby and Callaghan 2002, 2004b). The picture is much the same for the interannual variance of temperature during this season. Together, anomalous EP flux and equatorial wind account for as much interannual variance during late winter, as does anomalous EP flux alone during early winter.

Two parallel integrations are performed. In one, tropospheric wave structure is amplified to intensify by 20% and, simultaneously, the QBO is imposed in its easterly phase. These conditions correspond to anomalously high ozone in the observational record (Salby and Callaghan 2002, 2004b). In the second integration, tropospheric wave structure is suppressed to weaken by 20% and, simultaneously, the QBO is imposed in its westerly phase. Those conditions correspond to anomalously low ozone in the observational record. Differencing the integrations then leads to the climate sensitivities of temperature and ozone: the anomalous structures introduced by a one-standard-deviation increase in anomalous forcing of the residual circulation. As in section 3a, those structures represent the computational analog of anomalous structures that were composited from the observed record.

Figure 6 plots the anomalous wintertime tendency of temperature during late winter. The tendency ΔTMar–Nov has the same general form as prevails during early winter (Fig. 2); however, the strong positive anomaly over the Arctic now maximizes in the lower stratosphere, where it exceeds 10 K per season. Reflecting a gradual descent and intensification of anomalous temperature (Fig. 3), the same seasonal transition emerges from the observational record, wherein ΔTMar–Nov likewise approaches 10 K per season (Salby and Callaghan 2004b; see their Fig. 5).

Fig. 6.
Fig. 6.

Anomalous wintertime tendency of temperature (K per season) during late winter (November–March), forced by a one-standard-deviation intensification of EP flux from the troposphere and, simultaneously, an easterly swing of the QBO.

Citation: Journal of the Atmospheric Sciences 68, 7; 10.1175/2011JAS3671.1

Also distinguishing ΔTMar–Nov is anomalous structure in the tropics. Although comparatively weak, a cold anomaly appears in the equatorial lower stratosphere, capped overhead by a warm anomaly in the equatorial upper stratosphere. Through thermal wind balance, the cold anomaly reflects easterly shear that characterizes the QBO’s easterly phase. It also reflects anomalous upwelling and adiabatic cooling, associated with the QBO’s residual circulation. To either side of the cold equatorial anomaly are warm anomalies in the subtropics of each hemisphere. They mark anomalous downwelling, likewise associated with the QBO’s residual circulation. Note that this consequence of the QBO expands the polar anomaly equatorward from its extent in the absence of the QBO (Fig. 2). Previously found poleward of 50°N, the warm anomaly over the Arctic now extends equatorward of 30°N. Thermal wind balance then implies a commensurate change of the polar-night jet. Implied, in particular, is modified wind in the subtropics, where EP flux is absorbed inside the critical region of planetary waves.

Figure 7 plots the anomalous wintertime tendency of ozone mixing ratio. Anomalous is large and positive in the Arctic stratosphere. It reflects intensified downwelling of ozone-rich air, which is transferred poleward through isentropic mixing by planetary waves. The positive anomaly over the Arctic is associated with mixing ratio surfaces that, under conditions of intensified wave forcing, have been driven into coincidence with isentropic surfaces. A signature of isentropic mixing, the same structure is manifest in observed surfaces—during warm winters, but not during cold winters. Such mixing, which prevails at high latitude under conditions of intensified wave forcing, involves transient eddies that extend farther poleward, transferring ozone-rich air deeper into the Arctic (see, e.g., Gray et al. 2003). By homogenizing rO3, isentropic mixing fills the ozone void at high latitude that exists under conditions of weak wave forcing. The latter is manifest in observed ozone structure during cold winters, when surfaces at high latitude are deflected below isentropic surfaces, causing to decrease into the Arctic.

Fig. 7.
Fig. 7.

As in Fig. 6, but for the wintertime tendency of ozone mixing ratio (ppmv per season).

Citation: Journal of the Atmospheric Sciences 68, 7; 10.1175/2011JAS3671.1

The anomalous over the Arctic increases upward, approaching 2 ppmv per season near 10 mb. Compensating it at subpolar latitudes is negative ; however, the negative anomaly is largely confined to the middle and upper stratosphere. In the lowermost stratosphere, where total ozone is concentrated, anomalous remains positive across most of the winter hemisphere. In the equatorial middle stratosphere is another positive anomaly. It coincides with the cold anomaly of the QBO at upper levels, where rO comes under photochemical control. This suggests that the enrichment of ozone in the equatorial stratosphere follows from reduced destruction of O3 through recombination with O (see, e.g., Haigh and Pyle 1982). Although in the equatorial stratosphere approaches 1 ppmv per season, that anomaly has little support at lower levels, which determine ozone column abundance.

Plotted in Fig. 8 is the anomalous wintertime tendency of total ozone (solid). Anomalous Δ〈O3Mar–Nov is negative in the tropics and the subtropics of the summer hemisphere, where it has values of about −2 DU per season. From there, it increases northward until reaching a plateau at midlatitudes of the winter hemisphere. There, Δ〈O3Mar–Nov has values of about 4 DU per season. Similar structure characterizes the observed climate sensitivity of total ozone that was composited from the TOMS record, albeit with values in the plateau of about 8 DU per season (Salby and Callaghan 2002, their Fig. 16).

Fig. 8.
Fig. 8.

As in Fig. 6, but for the wintertime tendency of total ozone (DU per season) in the presence of homogeneous/gas-phase chemistry (solid) and inclusive of heterogeneous processes (dashed).

Citation: Journal of the Atmospheric Sciences 68, 7; 10.1175/2011JAS3671.1

In the observational record, the plateau at midlatitudes is associated with the QBO. When Δ〈O3Mar–Nov is composited against alone, it disappears. The modeled climate sensitivity has the same dependence: When derived from parallel integrations exclusive of the QBO, the plateau at midlatitudes is absent (not shown).

Poleward of 60°N, anomalous Δ〈O3Mar–Nov in Fig. 8 increases sharply, exceeding 24 DU per season over the Arctic. This may be compared against the observed anomaly over the Arctic, wherein Δ〈O3Mar–Nov approaches 28 DU per season.4 The computed structure of anomalous ozone reproduces the salient features of the anomaly that was composited from the observed record, albeit with values slightly lower over the Arctic.

Superimposed in Fig. 8 is the anomalous wintertime tendency of total ozone under analogous conditions, but now inclusive of heterogeneous processes that activate chlorine and accelerate ozone destruction (dashed). Equatorward of 50°S, the contribution to anomalous Δ〈O3Mar–Nov from heterogeneous chemistry is less than 0.5 DU per season. At higher latitude, it increases sharply, exceeding 6 DU per season over the winter pole. The total anomaly then exceeds 30 DU per season. This is somewhat greater than the observed anomaly (~28 DU per season). However, in view of observational uncertainties over the Arctic, the computed anomaly is rather close.

Most of the ozone anomaly introduced through heterogeneous processes enters under cold conditions (anomalously weak forcing of residual motion). Temperature over the Arctic then attains a minimum of 187 K. The anomalously cold temperature supports increased NAT and swelling of sulfate aerosol. Increased surface area then accelerates chlorine activation and ozone destruction. Under warm conditions (anomalously strong forcing of residual motion), minimum temperature over the Arctic is some 10 K warmer. The strong temperature dependence of heterogeneous processes then limits additional activation of chlorine. Relative to warm winters, Arctic ozone during cold winters therefore becomes anomalously sparse. The preponderance of anomalous ozone, however, follows from other sources.

Of the 30 DU per season over the Arctic, heterogeneous chemistry accounts for about 20%. This is well short of the contribution to anomalous ozone that has hitherto been ascribed to chemical depletion, on the premise of air over the Arctic being isolated (e.g., Müller et al. 2002; Tilmes et al. 2003; Rex et al. 2004).5 However, the computed contribution from heterogeneous chemistry is consistent with the contribution derived from the observed record of the Solar Backscatter Ultraviolet, version 8 (SBUV-V8) instrument, in light of differences of ozone structure between warm and cold winters (Salby and Callaghan 2007). It is also consistent with CTM integrations, which reproduce observed changes of Northern Hemisphere ozone exclusive of heterogeneous processes (Hadjinicolaou et al. 2002). The remaining approximately 80% of anomalous ozone between warm and cold winters is introduced by anomalous transport. It follows from anomalous downwelling of ozone-rich air, which is transferred into the Arctic through isentropic mixing by planetary waves. Observed ozone structure, wherein mixing ratio surfaces are driven into coincidence with isentropic surfaces during warm winters, leads to a similar conclusion.

The computed climate sensitivities of temperature and ozone reproduce the major features of those composited from the observational records of ECMWF and TOMS. Notice that anomalous temperature structure in Fig. 6 resembles that during early winter (Fig. 2), when it follows exclusively from anomalous . In view of the resemblance, it is instructive to isolate the contribution from the QBO.

Parallel integrations are now performed with fixed, but with the QBO in extremal phases. Plotted in Fig. 9 is the anomalous wintertime tendency of temperature that results. In the tropics, anomalous ΔTMar–Nov is close to that recovered earlier, when anomalous and the QBO are both represented (Fig. 6). The cold anomaly over the equator and accompanying warm anomalies in the subtropics reflect the residual circulation of the QBO. Consistent with this interpretation is the distribution of anomalous residual mean motion (not shown). Here, is characterized by anomalous upwelling over the equator, flanked in the subtropics by anomalous downwelling.

Fig. 9.
Fig. 9.

As in Fig. 6, but for the QBO contribution alone (EP flux fixed).

Citation: Journal of the Atmospheric Sciences 68, 7; 10.1175/2011JAS3671.1

At polar latitudes, the anomalous structures in Figs. 6 and 9 differ conspicuously. Although having form similar to that derived inclusive of anomalous , anomalous ΔTMar–Nov over the Arctic is now small. Less than 3 K per season, the anomalous tendency of temperature is broadly consistent with the observed wintertime anomaly associated with the QBO (Naito and Yoden 2005). It is conveyed into the Arctic through planetary waves, which interact with uEQ to drive an anomalous residual circulation. Although it extends to the pole, anomalous downwelling is much weaker than that induced by anomalous .

The corresponding tendency of total ozone is plotted in Fig. 10 (solid). In the tropics, anomalous structure is again close to that recovered when anomalous and the QBO are both included (Fig. 8). At extratropical latitudes, however, anomalous Δ〈O3Mar–Nov increases to only about 7 DU per season, far short of the 24 DU per season that is recovered in the presence of both mechanisms that force anomalous residual motion. The extratropical maximum is found at midlatitude, near the plateau in Fig. 8. A comparison of values indicates that anomalous total ozone is controlled by the QBO equatorward of 40°N, but by anomalous at higher latitudes.

Fig. 10.
Fig. 10.

As in Fig. 9, but for the wintertime tendency of total ozone.

Citation: Journal of the Atmospheric Sciences 68, 7; 10.1175/2011JAS3671.1

Superimposed in Fig. 10 is the anomalous wintertime tendency of 〈O3〉 under analogous conditions, but inclusive of heterogeneous processes (dashed). As in the presence of anomalous , the contribution from heterogeneous chemistry is small equatorward of 50°N. It increases at higher latitude. However, because anomalous temperature over the Arctic is weak, so too is the contribution from heterogeneous chemistry, now adding only about 1.5 DU per season. The anomalous wintertime tendency of 〈O3〉 there approaches 6 DU per season. The overall picture, however, is unchanged: At high latitude, where anomalous total ozone in Fig. 8 is large, most of the anomaly is introduced by anomalous . This resonates with the observational picture, wherein most of the interannual variance is carried by anomalous (Salby and Callaghan 2002). Likewise, of the calculated anomaly in Arctic ozone, some 80% is introduced by anomalous transport. A similar conclusion follows from the observational record of SBUV-V8.

4. Seasonality of anomalous temperature and ozone

Anomalous temperature and ozone at high latitude are each accompanied at low latitude by anomalous structure of opposite sign. It is instructive to consider how those compensating anomalies evolve during the course of the year.

The seasonality of anomalous temperature is considered in its column average 〈T〉 above θ = 360 K, which serves as an analog of the column abundance of ozone, 〈O3〉. The seasonal-mean tendency is then evaluated in terms of a sliding 3-month difference,
e2
centered at month Mj. The anomalous seasonal tendency is averaged over polar latitudes, poleward of 60°N, and over subpolar latitudes, from 60°N to 30°S.

Figure 11 plots the anomalous seasonal-mean tendency of temperature, as a function of central month. At high latitude (solid), is large and positive near winter solstice, when anomalous forcing of the residual circulation is strong. It decreases during late winter, mirroring the seasonal decline of . The anomalous tendency accumulates during the disturbed season, magnifying anomalous temperature at high latitude. Then, following spring equinox and the final warming, the anomalous temperature tendency reverses: becomes negative. The negative tendency gradually erases the positive temperature anomaly that developed during the disturbed season from the positive anomalous tendency. The anomalous tendency remains negative until midsummer, by which time and anomalous Arctic temperature have decayed to small values.

Fig. 11.
Fig. 11.

Anomalous seasonal-mean tendency of column-averaged temperature above θ = 360 K, as function of central month, averaged poleward of 60°N (solid) and from 60°N to 30°S (dashed). Column average corresponds to vertical integral over pressure, making it analogous to total ozone. Seasonal-mean tendency calculated as sliding 3-month difference.

Citation: Journal of the Atmospheric Sciences 68, 7; 10.1175/2011JAS3671.1

At low latitude (dashed), Δ〈TMar–Nov exhibits similar seasonality but is smaller (cf. Fig. 2). It is just opposite to the seasonality at high latitude. During winter, when the anomalous tendency at high latitude is positive, the anomalous tendency at low latitude is negative. Then, when the anomalous tendency at high latitude reverses, so does the anomalous tendency at low latitude. It then becomes positive. Analogous to behavior at high latitude, the anomalous tendency remains positive until midsummer, by which time and the attending temperature anomaly have decayed to small values. The computed seasonality in Fig. 11, which is coherent but out of phase between polar and subpolar latitudes, mirrors anomalous temperature in the observed record from ECMWF (Salby and Callaghan 2004a).

The reversed temperature tendency that develops after the final warming in each region reflects radiative relaxation. Prevailing after the disturbed season, it gradually erases the positive anomaly at high latitude and the negative anomaly at low latitude that developed earlier from anomalous tendencies of opposite sign. This restoring influence drives thermal structure back toward climatological-mean temperature, after the perturbing influence of wintertime wave activity has disappeared.

Plotted in Fig. 12 is the anomalous tendency of total ozone. It undergoes similar seasonality. At high latitude (solid), is positive during early winter. It actually weakens then slightly, reflecting the incorporation of ozone-lean air over the Arctic from the mesosphere. During late winter, the anomalous ozone tendency increases, reflecting downwelling of ozone-rich air at subpolar latitudes that is subsequently transferred poleward via isentropic mixing by planetary waves. It is, in fact, such anomalous transport that distinguishes warm and cold winters in the observed record (Salby and Callaghan 2007). In the computed behavior, isentropic mixing into the Arctic increases anomalous ozone until shortly after spring equinox. Then, during April, the anomalous ozone tendency reverses: it becomes negative. The negative tendency gradually erases the positive ozone anomaly that developed earlier from the positive anomalous tendency. The anomalous tendency remains negative until midsummer, by which time and anomalous Arctic ozone have decayed to small values.

Fig. 12.
Fig. 12.

As in Fig. 11, but for the seasonal-mean tendency of total ozone.

Citation: Journal of the Atmospheric Sciences 68, 7; 10.1175/2011JAS3671.1

At low latitudes (dashed), the anomalous ozone tendency exhibits similar seasonality but is smaller. However, like anomalous temperature, it is just opposite to the seasonality at high latitude. During winter, when the anomalous tendency at high latitude is positive, the anomalous tendency at low latitude is negative. (Notice that during early winter both weaken slightly.) Then, when the anomalous tendency at high latitude reverses, so does the anomalous tendency at low latitude. Analogous to behavior at high latitude, the anomalous tendency at low latitude remains positive until midsummer, by which time and the attending ozone anomaly have decayed to small values. The computed seasonality in Fig. 12—which, like anomalous temperature, is coherent but out of phase between polar and subpolar latitudes—mirrors that of anomalous ozone in the observed record from TOMS (Salby and Callaghan 2004a).

The reversed ozone tendency that develops in each region after the final warming reflects photochemical relaxation. It gradually erases the positive anomaly at high latitude and the negative anomaly at low latitude that developed earlier from anomalous tendencies of opposite sign. This restoring influence drives ozone back toward climatological-mean levels, after the perturbing influence of wintertime wave activity has disappeared.

By summer’s end, the anomalous tendencies of temperature and ozone have each been reduced to small values, at high latitude and at low latitude. Little memory of the disturbed season then remains. Much the same picture follows from the observational record (Hadjinicolaou et al. 1997; Salby and Callaghan 2002).

The calculated anomalies in Figs. 11 and 12 reproduce the observed seasonality of anomalous temperature and ozone. Evolving out of phase between high and low latitudes, those anomalies represent random changes between years—changes that are coupled to anomalous forcing of the residual circulation (see section 3). Very similar seasonality characterizes systematic changes, which comprise observed trends of temperature and ozone (Salby and Callaghan 2004a).

5. Discussion and conclusions

The numerical integrations recover the salient features of the observed climate sensitivities of temperature and ozone, which have been composited from the ECMWF and TOMS records. Characterized by a strong anomaly of one sign at polar latitudes and a comparatively weak anomaly of opposite sign at subpolar latitudes, each bears the signature of the residual mean circulation. In concert with isentropic mixing by planetary waves, residual mean transport sharply increases temperature and ozone over the Arctic.

During early winter, anomalous temperature and ozone are accounted for almost entirely by anomalous EP flux from the troposphere. The same is true in the observed record. The computed climate sensitivity then reproduces major features of the observed structure. During late winter, anomalous EP flux and equatorial wind are both required to account for the observed anomalies.

The results are consistent with observational findings that, jointly, anomalous EP flux from the troposphere and the QBO are the major mechanisms behind interannual changes in the stratosphere. In the observed record, the lion’s share of the interannual variance is carried by anomalous alone. The same feature characterizes the numerical integrations. Nevertheless, the two mechanisms interfere, especially during late winter and spring. This may explain why changes of polar ozone track the QBO during some years, but not during others (Garcia and Solomon 1987; Salby and Callaghan 2002, their Fig. 3). Including both mechanisms reproduces the observed response of temperature and ozone, accounting for nearly all of the interannual variance over the Northern Hemisphere.

Integrations distinguished by the inclusion of heterogeneous processes isolate the respective chemical contribution to anomalous ozone. Coupled to anomalous temperature, that contribution, together with anomalous transport, is introduced by anomalous forcing of the residual circulation. The contribution from heterogeneous chemistry is greatest over the Arctic, where anomalous temperature induced by anomalous downwelling is large. There, heterogeneous chemistry accounts for about 20% of the computed anomaly of total ozone. The remaining approximately 80% of the computed anomaly follows from anomalous transport of ozone-rich air, through anomalous downwelling and isentropic mixing by planetary waves. A similar conclusion follows from observed ozone structure, which is distinguished between warm and cold winters by its relationship to isentropic surfaces (Salby and Callaghan 2007).

The climate sensitivity of temperature has structure very similar to that associated with the Arctic Oscillation (Thompson and Wallace 2000). Out of phase between high and low latitudes, each form of interannual variability bears the signature of the residual mean circulation. The numerical integrations, in which anomalous temperature is formally linked to residual motion, make its involvement explicit.

The anomalous residual circulation couples the stratosphere to the troposphere through transfers of mass. Through downwelling over the Arctic, stratospheric air is transferred into the troposphere. Air must be returned to the stratosphere at the same rate, through upwelling of tropospheric air at lower latitudes. Changes over the Arctic should therefore be compensated by opposing changes at the tropical tropopause. Shaped by deep convection, the tropical tropopause is, in turn, coupled to the Hadley circulation. Observed changes in the Arctic stratosphere, in fact, vary coherently with changes in the tropical upper troposphere (Salby and Callaghan 2005). They reflect a global interaction between the Brewer–Dobson circulation and the Hadley circulation.

The climate sensitivities computed here represent the response of stratospheric temperature and ozone to random changes between years. The latter enter through changes of momentum transmission from the troposphere by planetary waves and of equatorial wind associated with the QBO. Both introduce anomalous residual motion. The corresponding response of temperature and ozone has a particular structure, as well as a particular seasonality. Each bears the signature of the residual mean circulation. Having much the same structure and seasonality are systematic changes in the observational record, which comprise trends of stratospheric temperature and ozone. How the residual circulation enters systematic changes remains to be determined. Nevertheless, the strong resemblance of systematic changes to random changes, wherein the residual circulation figures centrally, implies its involvement in both.

Acknowledgments

The author is grateful for constructive comments provided during review and to D. Kinnison and S. Walters for guidance with the implementation of MOZART-3.

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1

Day-to-day differences in structure can also develop internally, for example, through local instability that is introduced by planetary waves (e.g., Scaife and James 2000; Gray et al. 2003). However, such differences are overshadowed by those that follow from coherent differences in , which, in the observed record, account for the preponderance of interannual variance (Salby and Callaghan 2002).

2

Observed changes of EP flux and the QBO are augmented by the solar cycle. By modifying ozone heating in the upper stratosphere, it provides a third source of interannual variability. The solar cycle can modulate the QBO’s influence, and even the QBO itself (Labitzke and van Loon 1988; Naito and Hirota 1997; Gray et al. 2004; Salby and Callaghan 2006). However, it can exert little influence on behavior that has been averaged over several decades, against which these integrations are compared.

3

Changes of wind structure in the stratosphere can, through reflection, occasionally influence upward EP flux from the troposphere. However, on seasonal-mean time scales, such influence must be secondary, as, in the observed record, EP flux prescribed from tropospheric structure accounts for the preponderance of interannual variance.

4

The accuracy of observed structure over the Arctic is limited by restricted coverage of TOMS during November, when high latitudes are shrouded in darkness.

5

Of order 60 DU (deviations of ±30 DU), prior estimates of the heterogeneous contribution are as large as the rms difference between warm and cold winters, 56 DU (deviations of ±28 DU).

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