Abstract During the Cretaceous (145.5-65.5 Ma; Gradstein et al. 2004 ). Central Europe was part of the European continental plate, which was bordered by the North Atlantic ocean and the Arctic Sea to the NW and north, the Bay of Biscay to the SW, the northern branch of the Tethys Ocean to the south, and by the East European Platform to the east ( Fig. 15.1 ). The evolution of sedimentary basins was influenced by the interplay of two main global processes: plate tectonics and eustatic sea-level change. Plate tectonic reconfigurations resulted in the widening of the Central Atlantic, and the opening of the Bay of Biscay. The South Atlantic opening caused a counter-clockwise rotation of Africa, which was coeval with the closure of the Tethys Ocean. Both motions terminated the Permian-Early Cretaceous North Sea rifting and placed Europe in a transtensional stress field. The long-term eustatic sea-level rise resulted in the highest sea level during Phanerozoic times ( haq et al. 1988;Hardenbol et al. 1998 ). Large epicontinental shelf areas were flooded as a consequence of elevated spreading rates of mid-ocean ridges and intra-oceanic plateau volcanism, causing the development of extended epicontinental shelf seas and shelf-sea basins ( Hays & pitman 1973 ; Larson 1991 ). A new and unique lithofacies type, the pelagic chalk, was deposited in distal parts of the individual basins. Chalk deposition commenced during middle Cenomanian-early Turanian times. Chalk consists almost exclusively of the remains of planktonic coccolithophorid algae and other pelagic organisms, and its great thickness reflects a high rate of production of the algal tests. The bulk of the grains are composed of lowmagnesium calcite, representing coccolith debris with a subordinate amount of foraminifers, calcispheres, small invertebrates and shell fragments of larger invertebrates ( Håkansson et al. 1974 ; Surlyk & Birkelund 1977 ; Nygaard et al. 1983 ; Hancock 1975 , 1993 ).

Abstract Compaction is one of the major processes by which sediments lose porosity and begin the transformation to sedimentary rocks. Compaction is driven mainly by overburden loading and involves changes in the packing density of constituent grains. This is accomplished initially through grain reorientation and repacking accompanied by water expulsion from porous sediments. With additional overburden loading, fracturing and cleavage of brittle grains and plastic deformation of ductile grains contribute to increased packing density and concomitant loss of pore space. Further reduction of intergranular pore space, beyond that produced by “mechanical compaction”, results from pressure-solution processes, sometimes termed “chemical compaction”. Chemical compaction includes selective dissolution and interpenetration at grain-to-grain contacts, as well as broader dissolution along solution seams and stylolites. Although more common in carbonate rocks, chemical compaction features are widespread in clastic terrigenous deposits as well (e.g., Heald, 1955; Walderhaug and Bjørkum, 2003).