Changes of total versus extreme precipitation and dry periods until the end of the twenty-first century: statistical assessments for the Mediterranean area

Abstract

Changes of total precipitation, extreme precipitation, and dry periods in the Mediterranean area until the end of the twenty-first century have been assessed by means of statistical downscaling. Generalized linear models using predictors describing the large-scale atmospheric circulation as well as thermodynamic conditions have been applied for the projections under A1B and B1 scenario assumptions. The results mostly point to reductions of total and extreme precipitation over the western and central-northern Mediterranean areas in summer and autumn and to increases in winter. In contrast, over the eastern Mediterranean area widespread precipitation increases are assessed in summer and autumn, whereas reductions dominate in winter. In spring, total and extreme precipitation decreases prevail over the whole Mediterranean area. Total and extreme precipitation decreases mostly come along with increases of the maximum dry period length. Vice versa precipitation increases are commonly accompanied by a shortening of the maximum dry period length.

Appendix 1

1.1 Calculation of convective indices

The Showalter Index (Showalter 1953) is used to describe convective instability and is calculated as follows:

$$ {\hbox{Showalter Index}} = {T_{{{5}00}}} - {\hbox{TLC}}{{\hbox{L}}_{{{5}00}}} $$

(1)

T ₅₀₀: temperature at the 500 hPa level

TLCL₅₀₀: temperature that an air parcel will achieve if it is lifted dry adiabatically from the 850 hPa level to its condensation level and then moist adiabatically to the 500 hPa level.

The difference between the 500 hPa temperature and the lifted air parcel temperature provides an estimate of the instability. It should be noted that the Showalter Index does not take into account the atmospheric conditions below the 850 hPa level. To calculate the Showalter Index from the reanalysis data and the GCM data, 850 hPa air temperature and 850 hPa relative humidity are used to calculate the dew point temperature, and from this the height of the condensation level at each grid box. The difference of the geopotential heights of the 500 and 850 hPa levels is used to obtain the absolute height of the air parcel lifting. The value of the moist adiabatic temperature gradient above the condensation level is readjusted in steps of 100 m uplift.

CIN, characterizing the presence of large- or small-scale lifting mechanisms, is represented by a proxy, following Myoung and Nielsen-Gammon (2010):

$$ {\hbox{CIN}} = T\left( {\hbox{inv}} \right)--{T_d}(s) $$

(2)

T _d (s): dewpoint temperature at the surface

T(inv): virtual temperature at some level just above the mixed layer or within a capping.

According to Myoung and Nielsen-Gammon (2010), surface pressure values can be taken to obtain the associated best proxy levels for the calculation of the proxy of CIN. In the present study, the assumed best proxy level for each grid box of the predictor domain is determined by taking Table 1 in Myoung and Nielsen-Gammon (2010). Thus, for the calculation of CIN from the reanalysis data and the GCM output, surface dewpoint temperatures are required as well as surface pressure data (to obtain the best proxy level), and additionally specific humidity and air temperature of the best proxy level (for the calculation of the virtual temperature).