The spectrum of geochemical compositions of Oceanic Island Basalts (OIBs) and their systematic differences from Mid-Ocean Ridge Basalts (MORBs) reveal that the Earth’s mantle is chemically and isotopically heterogeneous. Two main processes, both related to plate tectonics, contribute to the creation of mantle heterogeneities: (1) partial melting generates melts enriched in incompatible elements and leaves a depleted residual rock; and (2) subduction of the oceanic lithosphere injects heterogeneous material at depth, in particular, altered oceanic crust and continental/oceanic sediments. Moreover, delamination and foundering of metasomatized subcontinental lithospheric mantle might have been important in the early Earth history, when plate tectonics did not operate as today. The fate of the subducted plate is still a matter of debate; presumably some of it is stirred by convection and some may segregate at the base of the mantle, in particular the oceanic crust, which is compositionally denser than the pyrolitic mantle. The view of the lower mantle as a “graveyard” of subducted crust prevailed for decades and was supported by the Hofmann and White ( 1982 ) observation that the geochemical fingerprint of most OIB reveals the presence of ancient recycled crust. However, recent geochemical data on short-lived systems ( e.g. 182 Hf→ 182 W has a half-life of 8.9 My) showed that some hotspots, namely Hawaii, Samoa, Iceland and Galápagos, have a negative µ 182 W anomaly. This discovery prompted a change in our view of the deep mantle because anomalies in short-lived systems require additional processes, which include, but are not limited to, the preservation of ‘pockets’ of melt from a primordial magma ocean, and/or chemical reactions between the metallic core and the silicate mantle. Exchanges at the core-mantle boundary would cause a negative µ 182 W anomaly, and might also add 3 He to mantle material later entrained by plumes. It is now clear that some plumes probe the deepest mantle and are highly heterogeneous, as revealed by isotope ratios from long-lived radiogenic systems, noble gases and short-lived isotope systems. Here I will focus on the dynamics of plumes carrying compositional and rheological heterogeneities. This contribution attempts to be pedagogic and multi-disciplinary, spanning from seismology to geochemistry and geodynamics.

Collisions and impact processes have been important throughout the history of the solar system, including that of the Earth. Small bodies in the early solar system, the planetesimals, grew through collisions, ultimately forming the planets. The Earth started growing ca. 4.56 Ga in this way. Its early history was dominated by violent impacts and collisions, of which we only have circumstantial evidence. The Earth was still growing and had reached ∼70%–80% of its present mass when at ca. 4.5 Ga a Mars-sized protoplanet collided with Earth, leading to the formation of the moon—at least according to the currently most popular hypothesis of lunar origin. After its formation, the moon was subjected to intense post-accretionary bombardment between ca. 4.5 and 3.9 Ga. In addition, there is convincing evidence that the Moon experienced an interval of intense bombardment with a maximum at ca. 3.85 ± 0.05 Ga; subsequent mare plains as old as 3.7 or 3.8 Ga are preserved. It is evident that if a late heavy bombardment occurred on the Moon, the Earth must have been subjected to an impact flux at least as intense as that recorded on the Moon. The consequences for the Earth must have been devastating, although the exact consequences are the subject of debate (total remelting of the crust versus minimal effects on possibly emerging life forms). So far, no unequivocal record of a late heavy bombardment on the early Earth has been found. The earliest rocks on Earth date back to slightly after the end of the heavy bombardment, although there are relict zircons that have ages of up to 4.4 Ga (in which, however, no impact-characteristic shock features were found so far). In terms of evidence for impact on Earth, the first solid evidence exists in the form of various spherule layers found in South Africa and Australia with ages between ca. 3.4 and 2.5 Ga; these layers represent several (the exact number is still unknown) large-scale impact events. The oldest documented (and preserved) impact craters on Earth have ages of 2.02 and 1.86 Ga. Thus, the impact record for more than half of the geological history of the Earth is incomplete and not well preserved, and we mostly have only indirect evidence regarding the impact record and its effects during the first 2.5 b.y. of Earth history.