Some mechanisms of mid-Holocene climate change in Europe, inferred from comparing PMIP models to data

Abstract

We propose a new approach for comparing mid-Holocene climates from 18 PMIP simulations with climate reconstructions of winter and growing season temperatures and the annual water budget inferred from European pollen and lake-level data. A cluster analysis is used to extract patterns of multivariate climate response from the reconstructions; these are then compared to the patterns simulated by models. According to paleodata, summers during mid-Holocene were warmer-than-present in the north, and cooler-than-present in the south, while winters were colder-than-present in the southwest but milder-than-present in the northeast. Whereas warmer summers and colder winters may easily be explained as a direct response to the amplified seasonal cycle of insolation during the mid-Holocene, the other recorded responses are less straightforward to explain. We have identified, from the models that correctly simulate the recorded climate change, two important atmospheric and hydrological processes that can compensate for direct insolation effects. First, a stronger-than-present airflow from southwestern Europe that veers to the north over Eastern Europe, in winter, can consistently explain the reconstructed changes in this season’s temperatures and water budget. Second, the increased winter soil moisture allows a shift of the partitioning of net radiative energy towards latent rather than sensible heat fluxes, thereby decreasing surface temperature during the following summer season. Our approach therefore solves one of the recurring problems in model-data comparisons that arises when a model simulates the correct response but in the wrong location (as a consequence, for instance, of model resolution and topography).

Appendix

We can estimate over Southern Europe (1) the temperature change in winter due to the insolation forcing alone; (2) the change in soil moisture required to overcome the effect of insolation change on summer temperature. At the surface, and at equilibrium:

$$R_{n} = SW + LW_{d} - \sigma T^{4}_{s} = SH + LH + G$$

(6)

where R _n is the net amount of energy absorbed by the surface and available for transfer into other energy forms, SW and LW _d are the solar and IR energy fluxes absorbed by the surface, SH,LH and G are respectively the sensible, latent and ground heat fluxes, T _s is the surface temperature, and σ the Stefan-Boltzmann constant.

Equation (6) can be rewritten under the following form (Bonan 1998):

$$S{\left( {1 - \alpha } \right)} + LW_{d} - \sigma T^{4}_{s} = K_{0} {\left[ {T_{s} - T_{r} } \right]} + \beta K_{1} {\left[ {e^{*} (T_{s} ) - e_{r} } \right]} + G$$

(7)

with

$$K_{0} = \frac{{\rho Cp}}{{r_{a} }}$$

(8)

and

$$K_{1} = \frac{{\rho Cp}}{{\gamma (r_{a} + r_{c} )}}$$

(9)

where α is the planetary albedo, S the incident solar radiation, r _a and r _c are the aerodynamic and canopy resistances, γ is a constant equal to 65.17, ρ is the density of air, Cp the specific heat at constant pressure, e*(T _s ) is the saturation vapor pressure at temperature T _s T _r and e _r are the temperature and the vapor pressure within the lower layer of the atmosphere, and β is the moisture availability term, varying between 0 (when the soil is dry) and 1 (when the soil is saturated). We can assume that G _6ky = G _0ky and that cloudiness and the circulation did not change LW _d6ky = LW _d0ky.

• 1. Assuming that atmospheric circulation has not changed dramatically, the gradients of temperature (T _s -T _r ) and humidity (e*(T _s )–e _r ) should remain constant between the present and the mid-Holocene climates. Assuming also that the moisture availability in the soil β is constant, we can write:

  $$\Delta T_{s} = {\left[ {\frac{{(1 - \alpha )}}{\sigma }} \right]}^{{1/4}} \cdot {\left[ {{\left( {S_{{6ka}} } \right)}^{{1/4}} - {\left( {S_{{0ka}} } \right)}^{{1/4}} } \right]}$$

  (10)

  Consequently, for α = 30%, S _0ky = 150W/m², S _6 ky = 140 W/m², a change of –10 W/m² in insolation change leads to a change of –3.5 °C in surface temperature.

• 2. If an increase of moisture availability β offsets the insolation effect, we can have: T _s6ky = T _s0ky . Assuming that atmospheric circulation has not changed during the summer, we can write:

  $$\Delta \beta = \frac{{\Delta S{\left( {1 - \alpha } \right)}}}{{{\left[ {e^{*} (T_{s} ) - e_{r} } \right]}.K_{1} }}$$

  (11)

So, for T _s = 25 °C, e*(T _s ) = 3166 Pa, we can assume reasonably that e _r = 0.7 · e*(T _s ) = 2216 Pa, r _a = 20 s/m, r _c = 430 m/s. In this case Δβ = 0.018165ΔS which means that a 20 W/m² increase in incident solar radiation can be offset by an increase 0.36 (no unit) in moisture availability β (ranging between 0 and 1). For a field capacity in this soil of 300 mm, an increase of 0.36 for β corresponds to a 108 mm increase in soil moisture.