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■ Fig. 3.2.2. A Depth of acidification with acid rain (FBW 1989; Last and Watling 1991) showing that when strong acids are deposited, horizons from A- to the upper C-horizon can be acidified, without becoming visible as soil horizon. The base saturation rises in the lower C-horizons. Due to cation uptake by roots and shedding of needles the organic matter on the forest floor also becomes more alkaline. B Changes in pH of the soil solution with continuing weathering as a consequence of the cumulative proton stress, i.e. acid deposition (after Schulze and Ulrich 1991). C Availability of nutrients dependent on the pH of the soil solution. (Larcher 1994)

creases only near the original bed rock. Stronger alkaline saturation and higher pH also occur in the deposited humus, as cations from roots from lower soil layers are taken up into the plant and reach the organic deposits via the litter.

The rate of soil acidification depends on the mineral constitution of the bed rock and the cu mulative acid stress (Fig. 3.2.2 B). On limestone soils with high CaC03 the imbalance of cations triggered by fixation of cations in organic matter or by relocation of cations is, at first, balanced by weathering of carbonate.

where

The released Ca ions occupy the charges that become free on the soil exchangers (see Chap. 2.3.1). With the loss of organic matter, e.g. by harvesting and leaching of cations from the clay minerals and humus complexes, CaC03 is continually consumed. Over a long period, the soil pH decreases and a reversible exchange of cations occurs with clay minerals and humus. With continuing loss of cations and H+-buffer-ing by metal oxides and hydroxides, a pH-de-pendent increase in the availability of certain metal ions occurs. Thus Mn becomes mobile at a pH between 5 and 4.2. At a pH of 4.2, the soil reaches another stable buffer system as with the availability of lime, but in this case the pH is buffered by A1 hydroxides (see Chap. 1.8); Fe buffers below pH 3.8. The availability of ions is very variable during the course of this process and each element is specifically dependent on the pH of the soil solution (Fig. 3.2.2 C; Larcher 1994).

The chemical changes in the soil are reversible, i.e. by fertilisation or liming, provided that the clay minerals are not restructured. As soon as the crystalline structure of silicates and clay minerals is changed (e.g. by dissolving the A1 lattice as a replacement of alkaline cations by protons, Chap. 2.1), a reverse into the original state is no longer possible, not even by supplying cations (liming of forests; see Chap. 3.5.1).

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