## 1. Introduction

[2] Stratospheric water vapor plays an important role for the radiation budget of the troposphere [*Held and Soden*, 2000] as well as the stratosphere [*Forster and Shine*, 2002] and for stratospheric chemistry, in particular for ozone [*Evans et al.*, 1998; *Dvortsov and Solomon*, 2001]. Water vapor enters the stratospheric “overworld” (the part of the stratosphere above 380 K potential temperature [*Holton et al.*, 1995]) through troposphere-to-stratosphere transport of air primarily in the tropics. It is widely agreed that the water vapor mixing ratio of air entering the stratosphere (henceforth [H_{2}O]_{e}) is limited to a few parts per million by volume (ppmv) by the extremely low temperatures at the tropical tropopause, but the details of the processes that control [H_{2}O]_{e} remain controversial [*Newell and Gould-Stewart*, 1981; *Danielsen*, 1993; *Sherwood and Dessler*, 2000; *Holton and Gettelman*, 2001]. Recently, *Fueglistaler et al.* [2005] have shown that annual mean and seasonal cycle of [H_{2}O]_{e} can be explained to within ∼0.2 ppmv by synoptic-scale temperature and winds, provided the full four-dimensional temperature history of air entering the stratosphere is taken into account.

[3] In addition to the seasonal cycle, there is significant interannual variability of stratospheric water vapor [e.g., *Randel et al.*, 2004], and there is also some evidence for a long-term positive trend [*Oltmans et al.*, 2000; *Rosenlof et al.*, 2001]. In this paper we present an analysis of interannual to multidecadal variability of stratospheric water vapor concentrations, and of their governing processes. We use the trajectory calculations of tropical troposphere-to-stratosphere transport (TST) of *Fueglistaler et al.* [2005] for the period 1979–2002 to predict [H_{2}O]_{e}, and analyze the processes controlling interannual variability in the model calculations. Systematic differences between model calculations and observations may be taken as indicative of processes affecting [H_{2}O]_{e} not taken into account in the model calculations.

[4] To focus ideas, let *F* = *F*_{1} + *F*_{2} be the total mass flux into the tropical stratosphere, and *F*_{1} be the mass flux resolved by the model calculations, whereas *F*_{2} is the mass flux due to unresolved processes. Then the total water flux is *F*_{1} · χ_{1} + *F*_{2} · χ_{2}, where χ_{1} is the water vapor mixing ratio predicted by the model calculations, and χ_{2} is the (unknown) water vapor mixing ratio of the unresolved flux. This decomposition allows one to account for systematic differences between χ_{1} and χ_{2}, and the (true) entry mixing ratio may be written as

where *f*_{1} ≡ *F*_{1}/*F* is the fraction of TST that is resolved, and *f*_{2} = 1 − *f*_{1} is the fraction not resolved by the model calculations. The trajectory calculations used here are based on wind and temperature fields of the reanalysis project, 40-year European Reanalysis (ERA-40) [*Simmons and Gibson*, 2000], of the European Centre for Medium-Range Weather Forecasts (ECMWF). Thus *F*_{1} is the mass flux into the tropical stratosphere as represented in ERA-40, and *F*_{2} is the flux due to processes not or poorly resolved in ERA-40, for example individual convective cells penetrating the tropopause. The model calculations then further assume that tropical cirrus clouds typically dehydrate the air to the saturation mixing ratio (χ^{sat}) of the synoptic-scale temperature field, and consequently the mixing ratio of tropical TST upon entry into the stratosphere is determined by the minimum saturation mixing ratio encountered during ascent, i.e.,

where the synoptic temperature *T*_{syn} is the temperature field as resolved by ERA-40. The model calculations employed here deliberately simplify many aspects of cloud microphysics and mesoscale dynamics (that affect ε in the equation above), and reflect essentially synoptic-scale dynamics and temperature fields in the tropical upper troposphere and lower stratosphere. For these calculations *Fueglistaler et al.* [2005] determined the total residual due to unresolved transport and simplifications *f*_{1} · ε + *f*_{2} · χ_{2} ≃ 0.2 ppmv, or about 5% of [H_{2}O]_{e}. It may be that one reason that the residual is small is that there are some cancelling effects between different neglected processes, in particular between incomplete sedimentation of cirrus clouds (that might imply positive ε) and the effect of mesoscale temperature fluctuations (e.g., due to gravity waves) that would tend to lower the minimum temperature compared to that of the synoptic-scale field (and hence might imply negative ε in the formulation above).

[5] In assuming *f*_{1} = 1 and ε = 0 the model calculations represent only interannual variability due to temperature variability and transport pathway variability of tropical TST as resolved by ERA-40. In particular the model calculations do not consider processes that could induce variability or longer-term changes in ε, or in the contribution to stratospheric water vapor from the unresolved flux (*f*_{2} · χ_{2}).

[6] The paper is organized as follows. Section 2 presents the data and the method to predict [H_{2}O]_{e}. Section 3 compares the model predictions for [H_{2}O]_{e} and water vapor mixing ratios in the lower tropical stratosphere to spaceborne observations beginning in 1986, and section 4 analyzes the processes governing interannual variability in the model calculations. Section 5 discusses the role of tropical zonal mean temperature anomalies (as opposed to localized temperature trends or transport anomalies) for [H_{2}O]_{e} anomalies. Employing a simple model for transport and chemistry in the stratosphere, we further compare in section 6 the model predictions to measurements in the midlatitude stratosphere. Section 7 recapitulates the results of the extensive comparisons presented in earlier sections, and discusses their implications for multidecadal trends. Section 8 finally provides a summary and conclusions.