Abstract Over the last 65 Ma, our world assumed its modern shape. This timespan is divided into the Palaeogene Period, lasting from 65 to 23 Ma and the Neogene, which extends up to the present day (see Gradstein & Ogg (2004) and Gregory et al. (2005) for discussion about the Quaternary). Throughout the Cenozoic Era, Africa was moving towards Eurasia in a northward direction and with a counterclockwise rotation. Numerous microplates in the Mediterranean area were compressed, gradually fusing, and Eurasia underwent a shift from a marine archipelago to continental environments, related to the rising Alpine mountain chains ( Figs 17.1 & 17.2 ). Around the Eocene-Oligocene boundary, Africa's movement and subduction beneath the European plate led to the final disintegration of the ancient Tethys Ocean. The Indo-Pacific Ocean came into existence in the east while various relict marine basins remained in the west. In addition to the emerging early Mediterranean Sea, another relict of the closure of the Tethys was the vast Eurasian Paratethys Sea. The Oligocene and Miocene deposits of Central Europe are largely related to the North Sea in the north, the Mediterranean Sea in the south and the intermediate Paratethys Sea and its late Miocene to Pliocene successor Lake Pannon. At its maximum extent, the Paratethys extended from the Rhône Basin in France towards Inner Asia. Subsequently, it was partitioned into a smaller western part consisting of the Western and the Central Paratethys and the larger Eastern Paratethys. The Western Paratethys comprises the Rhône Basin and the Alpine Foreland Basin of Switzerland, Bavaria and Austria. The Central Paratethys extends from the Vienna Basin in the west to the Carpathian Foreland in the east where it abuts the area of the Eastern Paratethys. Eurasian ecosystems and landscapes were impacted by a complex pattern of changing seaways and land bridges between the Paratethys, the North Sea and the Mediterranean as well as the western Indo-Pacific (e.g. Rögl 1998 ; Popov et al. 2004 ). This geodynamically controlled biogeographic differentiation necessitates the establishment of different chronostratigraphic/geochronologic scales. The geodynamic changes in landscapes and environments were further amplified by drastic climate changes during the Cenozoic. The warm Cretaceous climate continued into the early Palaeogene with a distinct optimum near the Palaeocene-Eocene boundary (Palaeocene-Eocene Thermal Maximum) and the Early Eocene (Early Eocene Climate Optimum). A gradual decrease in temperature during the later Eocene culminated in the formation of the first icesheets in Antarctica around the Eocene-Oligocene boundary ( Zachos et al. 2001 ; Prothero et al. 2003 ). A renewed warming trend that began during the Late Oligocene continued into the Middle Miocene with a climax at the Mid-Miocene Climatic Optimum. The turning point at around 14.2 Ma led to the onset of the Middle Miocene Climate Transition indicated by the cooling of surface waters and the expansion of the East Antarctic icesheet ( Shevenell et al. 2004 ). A final trend reversal during the Early Pliocene is reflected by a gentle warming until 3.2 Ma ( Zachos et al. 2001 ) when the onset of permanent Arctic glaciation heralded the Pleistocene ice ages (see Litt et al. 2008 ). The Cenozoic history of Central Europe is chronicled in a dense pattern of Palaeogene and Neogene basins. In addition to the more stable North Sea Basin, the majority of these basins were strongly influenced by the Alpine compressive tectonics which caused a general uplift of Europe during the Cenozoic (see Froitzheim et al. 2008 ; Reicherter et al. 2008 ). The marginal position of the seas covering the area and the considerable synsedimentary geodynamic control resulted in incomplete stratigraphic sequences with frequent unconformities, erosional surfaces and depositional gaps. This chapter deals with the Paleogene and Neogene (“Tertiary”) geological development of Central Europe and its adjacent areas. It is structured according to the main geological regions relevant for the Cenozoic: (1) The European Plate; (2) the Alps and Alpine Foredeep; (3) the Carpathians, their foredeep and the Pannonian Basins System; and (4) the Southern Alps and Dinarides. Each subchapter is arranged from west to east, and north to south.

Abstract A 160-m-long section measured in the lagoonal facies of the Middle Triassic Latemar platform (Dolomites, Italy) reveals a set of frequency components that we interpret as a strong Milankovitch signal. In this interpretation, all principal frequencies associated with the theoretical Middle Triassic precession index, P1 = 1/(21.7 ky), P2 = 1/(17.6 ky), and its modulations, E1 = 1/(400 ky), E2 = 1/(95 ky),and E3 = 1/(125 ky), were detected in a time-frequency evaluation of the cycles. A weak obliquity signal is also present in part ofthe section.Thus, the Latemar cycles appear to have recorded the clearest orbital forcing signal yet found in a carbonate platform. This astronomical calibration indicates that the section was deposited in ca. 3.1 My and therefore that the entire Latemar cyclic succession (~470 m) took at least 9 My to form. However, the calibration also leads to serious conflicts with other interpreted geological data: U/Pb radiometric ages of zircons collected from tuffites within theLatemar lagoon and in coeval basinal sediments point to a timescale that is five times shorter than this astronomically calibrated estimate; similar discrepancies arise when the average duration of Triassic ammonoid biozones or the sedimentation rates of coeval basinal series are considered. Nonetheless,all of the methods that have been used to estimate the time of formation of the Latemar platform continue to have shortcomings, and the contradictions among these different geologicalcalibrations remain unresolved.

Abstract A number of 3 rd -order depositional sequences have been observed in the Southern Alps, far more than previously known: of these, 6 develop principally in the Scythian, 4 in the Anisian, 3 in the Ladinian, 4 in the Carnian, at least 2 in the Norian and finally 2 in the Rhaetian. Lower Anisian to Lower Carnian depositional sequences are best dated by ammonoids, while in the Scythian ammonoids and conodonts are scanty. In the Upper Carnian, Norian and Rhaetian deposits of the Southern Alps, very rare ammonoids and conodonts are available. Therefore Scythian, Upper Carnian, Norian and Rhaetian depositional sequences have been defined on the basis of stratal patterns and the evolution of facies, being their chronostratigraphical position inferred from sections elsewhere in the world. The contemporaneous analysis of platforms and basins and of carbonate, mixed and siliciclastic deposits has made a good comprehension of facies migration possible. This approach was crucial in the definition of 3 rd -order depositional sequences. Moreover, on the basis of the previously defined 3 rd -order sequences and systems tracts, a number of 2 nd -order transgressive/regressive cycles have been pointed out. The sequence stratigraphic analysis compared with the tectonic history allowed the definition of the different phases of the basin evolution.