The earth is roughly 4600 Ma old (O Figure 16.1), and the oldest rocks currently recognized (Acasta gneiss from the Northwest Territories, Canada) are dated as 4030 Ma (Nelson 2004). The earliest Eon of earth history, the Hadean, thus left no directly observable documents. It is nowadays assumed that the initially high temperatures following the accretion of the earth dropped around 4000 Ma below 100° C. During the earliest Archean, the surface temperature was probably quite low (faint early sun). By that time, liquid water was present on the earth's surface, brought to the earth by comets (McClendon 1999). Yet the heavy bombardement with bodies exceeding 250 km lasted until 4200-3800 Ma. This must have led to repeated boiling of the oceans and the vaporization of the water (Nisbet and Sleep 2001, 2003).
Contrary to earlier beliefs (e.g., those assumed in the famous Miller experiment), most researchers today think that the Hadean and early U.S atmosphere was only mildly reducing, with mainly CO2 and N2, but also smaller amounts of CH4, NH3, H2 present (McClendon 1999; Raven and Skene 2003). Free oxygen builded perhaps through photodissociation of water vapor in the upper atmosphere but quickly reacted with Fe2+ and other unoxidized compounds. The early atmosphere was, therefore, devoid of free oxygen (McClendon 1999; Miller and Lazcano 2002).
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Under these conditions life originated (for reviews see Oró et al. 1990; Brack 1998; McClendon 1999; Fenchel 2002; Taylor 2005). Organic compounds could form in the atmosphere through UV-photolysis, electrical discharges, and major impact shocks (McClendon 1999; Lazcano 2001), and a major source of organic molecules was probably also comets and interplanetary dust (Lazcano 2001; Miller and Lazcano 2002), but early life certainly evolved in the sea (Raven and Skene 2003), and the first cells were probably heterotrophs (Lazcano and Miller 1996; but see Huber and Wachtershauser 1997; Wachtershauser 2000). As possible sites where cellular life evolved from prebiotic precursors, continental thermal springs, volcanic vents, warm hypersaline lagoons (Darwin's "warm little pond''), and deep submarine vents are the most likely candidates (Schopf 1999). In this respect, it is noteworthy that the earliest branches of both the Archea and Bacteria are thermophilic (Pace 1997). Yet an extraterrestrial origin of cellular life cannot totally be excluded (Horneck 2003).
Perhaps life originated sequentially several times but was always exterminated by the heavy meteorite bombardment ("impact frustration'') until life took hold around 3900 Ma (Schopf 2002). The "Universal Ancestor'' might have been a diverse community of cells that experienced extensive horizontal gene transfer (Woese 1998, 2002). Organismal lineages established themselves only with the subsequent splitting into the three domains "Bacteria," "Archea," and "Eukarya" (Woese 1998). The tree of life is thus at its base rather a web (Doolittle 1999) or even a ring of life (Rivera and Lake 2004).
The fossil record of the Archean is notoriously sparse. The presumed oldest cellular microfossils come from 3500 to 3450 Ma sediments in Western Australia (Schopf 1992a, 2004; but see Brasier et al. 2004). Almost as old (3400-3200 Ma) are microfossils in the metasediments of the Onverwacht and Fig tree groups (South Africa). From even older (3800 Ma) metamorphosed sedimentary rocks from Greenland (Isua Greenstone belt), putative microfossils (Isuasphaera) were described (Pflug 1978). Yet they are probably of inorganic origin (Bridgwater etal. 1981; Appel et al. 2003).
According to most authors (Mooers and Redfield 1996; Sheridan et al. 2003) the split between bacteria and archea-eukaryota must have occurred more than 3500 Ma (but see Doolittle et al. 1996). The Proterozoic fossil record (Hofmann and Schopf 1983; Mendelson and Schopf 1992; Schopf 1992a, b) documents an increasing diversification and complexity of cells but the evolution of metabolic pathways cannot be deduced from the morphology of the microfossils. Yet by 2700 Ma, oxygenic photosynthesis, methanogenesis, and methylotrophy had probably developed, perhaps also sulfate reduction and nitrogen fixation (Buick 2001).
In contrast to cellular fossils, the record of stromatolites is quite good. These are biosedimentary structures built by microbial mats which trap sediment particles (Riding 1991; Walter 2001). The oldest stromatolites date from 3500 to 3200 Ma sediments in Australia and South Africa, but their fossil record remains spotty until the Neoarchean (Walter 1983, 2001; Grotzinger and Knoll 1999). They diversified considerably in the Proterozoic, and in the Mesoproterozoic, a large number of different types had developed, growing sometimes to impressive sizes (Awramik and Sprinkle 1999; Walter 2001). Stromatolites declined both in diversity and abundance in the terminal Proterozoic, and in the Phanerozoic, they remained largely restricted to marginal marine environments. This decline was traditionally seen as a consequence of the rise of the metazoans, which consumed the microbial mats, but it might also have abiotic reasons (Awramik and Sprinkle 1999).
2450-2200 Ma, the earth witnessed the first well-documented ice age, the Huronian glaciation. As a consequence of insufficient time resolution, the exact duration of glacial intervals during this Paleoproterozoic time period is unknown (Young 2004). The onset and the termination seem to have been gradual (Young 2004). Perhaps an increase in atmospheric oxygen (see later) was the cause of the Huronian glaciation (Kopp et al. 2005). If methane was a major contributor to a greenhouse effect prior to that time, it would have been oxidized in a more oxygen-rich atmosphere, and the removal of this greenhouse gas could well lead to global cooling and the onset of an ice age.
At some time between 2500 and 1900 Ma, the atmosphere changed rather abruptly from reducing (pO2 < 1% present atmospheric level PAL) to oxidizing (pO2 > 15% PAL; O Figure 16.1), largely as a result of the activities of oxygenic photoautotrophs (Schopf 1992a; Holland 1994; for alternative scenarios of atmosphere evolution see Ohmoto 2004). Until about 2200 Ma, free oxygen was constantly removed by oxidation of weathered reduced minerals (Nisbeth and Sleep 2001; Lenton 2003). Detrital pyrite and uranite are common until 2200 Ma and indicate very low levels of free oxygen; otherwise these minerals would have been oxidized (Schopf 1992a; Holland 1994). Likewise, the genesis of banded iron formations (BIFs), which are absent in rocks younger than 1700 Ma, requires reducing conditions in the deeper parts of the seas (Simonson 2003).
The origin of the eukaryotes can be placed at some time before 2100 Ma. The oldest, currently recognized eukaryote is the 2100 Ma multicellular "algae" Grypania (Han and Runnegar 1992). Perhaps the origin of eukaryotes was linked to the increased oxygen content of the atmosphere and shallow waters where a nucleus and its protective membrane would be advantageous (Dyer and Obar 1994). Subsequent evolution of the eukaryotes took place by serial endosymbiosis (Margulis 1981) in which engulfed a-purple-bacteria became mitochondria and cyanobacteria became plastids (Dyer and Obar 1994; Pace 1997). In part, even secondary endosymbiosis must have occurred in which photosynthetic eukar-yotes were engulfed by nonphotosynthetic eukaryotes (Woese et al. 1990).
Little is known about the early evolution of the eukaryotes. Apart from the acritarchs that are a heterogeneous assemblage of planktonic, unicellular eukar-yotes known from the Late Paleoproterozoic and extending into the Phanerozoic (Martin 1993; Knoll 1994), other fossils of the eukaryotic clade include various carbonaceous films that are not easy to interpret (Hofmann 1994). Yet new molecular dates, which are in reasonable agreement with paleontological findings, give some clues as to when new groups originated. The divergence between protists and crown-group plants may be as young as 1100 Ma, the origin of the fungi is placed at 1000 Ma, the split between choanoflagellates and eumetazoans occurred at 900 Ma, and the first bilaterians might have made their appearance at 700 Ma (Douzery et al. 2004; see also Peterson et al. 2004; Peterson and Butterfield 2005). Yet these dates are not universally accepted, and previous molecular-clock estimates have yielded considerably older dates (see review in Erwin and Davidson 2002).
Between 800 and 600 Ma, the earth witnessed several large glaciations. The last two episodes are well dated. The Sturtian glaciation occurred around 710 Ma, and the Marinoan (Varanger) glaciation ended at 630-600 Ma (Allen and Hoffman 2005). Some, perhaps most, continental land masses had a near-equatorial position at that time. Glaciers might have reached equatorial regions and perhaps extended down to the sea level. According to the most dramatic scenario (Hoffman et al. 1998; Hoffman and Schrag 2002; but see Chandler and Sohl 2000; Poulsen 2003), temperatures initially dropped due to some unknown mechanism, but as soon as glaciers reached a critical extension in low latitudes, enough solar energy was reflected back into space that ice sheets could grow at an ever increasing rate. In this "runaway albedo'' model, not only all the land masses became ice covered, but the surface of the oceans were also globally frozen ("snowball earth''; Kirschvink 1992; Hoffman et al. 1998; Hoffman and Schrag 2002). The seas became anoxic and BIFs could accumulate again (Hoffman et al. 1998). Eventually, enough volcanic CO2 accumulated, and the glaciations ended abruptly. Later Neoproterozoic glaciations are documented, but these were not global in their extent (Knoll et al. 2004).
Shortly after the last of these major Neoproterozoic glaciations, the first metazoans enter the fossil record. Enigmatic soft-bodied fossils 610-542 Myr old are known from localities around the world and named Ediacara assemblages (after the Ediacara Hills in South Australia). Although the nature of the flattened, segmented, or quilted Ediacara organisms is still disputed (mainly cnidarians and annelids; Glaessner 1983, 1984; Jenkins 1992; or organisms not related to any of the extant animal phyla; Seilacher 1989, 1992; Buss and Seilacher 1994), they are accompanied by traces produced by bilaterian metazoans (Knoll and Carroll 1999; Valentine et al. 1999; Martin et al. 2000). Yet tracemakers were small and rare and did not disrupt the sediment. This only changed near the Precambrian-Cambrian boundary when bioturbation markedly increased and the sediment became unstable. This probably led to the extinction of the immobile Ediacara organisms (Seilacher 1999).
Within a short time after the beginning of the Cambrian at 542 Ma, the most remarkable episode in the history of life started, the so-called "Cambrian explosion.'' Within only 20 Myr, all animal phyla with preservable hard parts (with the exception of the bryozoans) appeared (Knoll and Carroll 1999; Erwin 2001a; Valentine 2002, 2004). With the beginning of the Cambrian and continuing throughout the Phanerozoic, we have a very reliable fossil record especially for the marine invertebrates with easily preservable hardparts. Fossiliferous localities that show exceptional preservation are interspersed in the stratigraphical record and provide us with information on the soft-bodied fauna (Bottjer et al. 2002). Information on the Phanerozoic history of life can be found in most textbooks on paleontology and historical geology (Cowen 2005; Stanley 2005), and extensive treatment is beyond the scope of that chapter. Details on selected episodes of radiation and extinction are given in the last part of this chapter.
During the Cambrian, life was exclusively marine and dominated by trilo-bites and a variety of other arthropods. Because the trilobites were highly diverse, are easily determined, and show a high species turnover, trilobites are the most important index fossils for this period. Brachiopods were small and belonged mostly to the inarticulate groups. Among the mollusks, hyoliths and monopla-cophorans were the most conspicuous ones, but in the Late Cambrian, the first small nautiloids appeared which marks the beginning of a highly successful group of marine predators. Cambrian echinoderms belonged mostly to groups that were immobile and became extinct during the early Paleozoic. Already during the Early Cambrian, the first true reefs built by the spongelike archeocyathans developed but archeocyathans became extinct at the end of the Early Cambrian, leaving a reef gap until the Middle Ordovician. The first chordates and agnathan fishes appear to have been rare with the exception of the conodonts.
The extinction events in the Late Cambrian affected most severely the trilobites and several echinoderm groups. Ordovician and Silurian seas became to be dominated by articulated brachiopods and stalked echinoderms (crinoids and blastoids). Although the first deep burrows appeared at that time, life was still mainly epibenthic. Large reefs dominated by tabulate and rugose corals and stromatoporoids developed, and bryozoans became an important component of marine hard bottoms. Among the planktonic organisms, the graptolites diversified considerably and have proven to be the most valuable index fossils for the Ordovician and Silurian. Large predators developed among the nektonic nautiloids and among the eurypterids. Several groups invaded freshwater environments, among them the arthropods and various fish groups. Colonializa-tion of the land started in the Silurian, first by plants (although algal crusts and fungi may have been present earlier), then by mites, arachnids, millipedes, and scorpions. Among the vertebrates, the appearance of the first jawed fishes in the Late Silurian was a major innovation.
During the Devonian and Carboniferous, land plants diversified in an explosive manner leading to the first true forests, and the first seed plants had developed by the Carboniferous. The first wingless insects appeared in the Early Devonian, but it was not until the Carboniferous that the first winged insects conquered the air. In the vertebrates, huge marine predators developed among the placoderms. The Devonian, also called the "age of fishes,'' saw the first appearance of the chondrichthyans and actinopterygians while the agnathans considerably diversified. The first tetrapods appeared during the Late Devonian, and the amniote egg evolved in the Late Carboniferous. Among the marine benthic organisms, tabulate-rugose-stromatoporoid reefs attained a new climax. Articulate brachiopods and stalked crinoids still dominated most seafloors. Among the nektonic organisms, the evolution of the ammonoids was a major innovation during the Early Devonian. This group provided the most important index fossils for the Devonian to Cretaceous periods. In the latest Paleozoic, life in the seas did not radically change, but the absence of the heavily armored fishes is notable and among the cephalopods, the stoutly shelled nautiloids also declined while the ammonoids flourished. Reef building was confined to smaller constructs after the Late-Devonian mass extinction. On land, the mammal-like reptiles were the dominant herbivores and carnivores.
After the end-Permian mass extinction, life in the seas and on land dramatically changed. Seafloors were no longer dominated by epifaunal brachiopods and crinoids but by gastropods and burrowing bivalves. Reef production came to a halt during the Early Triassic, and it was not until the Middle Triassic that scleractinian coral reefs became established. This type of reef building would dominate throughout the rest of the Meso- and Cenozoic. Only during part of the Cretaceous were the corals replaced as principal reef builders by a group of aberrant bivalves, the rudists. The open waters were dominated by ammonoids and actinopterygian fishes, and during the Mesozoic, three groups of planktonic organisms which play an eminent role in the biogeochemical cycles made their appearance: the planktonic foraminifers, the coccoliths, and the diatoms. The largest creatures of the seas and the top predators were the marine reptiles: nothosaurs, plesiosaurs, ichthyosaurs, marine crocodiles, turtles, and mosasaurs.
On land, the conifers, cycads, and ginkgos flourished and dominated the forests until the late Early Cretaceous, when they became increasingly replaced by the flowering plants (angiosperms). Among the amphibians, the last stegocephalians died out while the first frogs and salamanders appeared. After the end-Permian mass extinction, the mammal-like reptiles had become marginal players, and the archosaurs, most notably the dinosaurs, became the rulers of the earth. The pterosaurs were the first vertebrates to conquer the air by the Late Triassic, and in the Late Jurassic, the first birds evolved. Mammals arose already during the Late Triassic but remained mostly small and peripheral throughout the Mesozoic.
The mass extinction at the Cretaceous-Tertiary boundary, although much less severe than the extinction at the end of the Permian, nevertheless severely altered the structure of both marine and terrestrial ecosystems. In the seas, the shelled cephalopods became entirely extinct with the exception of a small group of nautilids, and from the marine reptiles only the turtles survived. The vacant ecospaces were filled up by actinopterygian fishes and mammals. The mammals also played a central role in the restructuring of the terrestrial communities. During the remarkable radiation in the Paleogene, mammals occupied almost all available niches on land, invaded the seas (whales, pinnipeds, sea cows), and conquered the air (bats). Birds equally radiated considerably, and in some regions large, flightless birds even became the top predators. After the Eocene-Oligocene climatic revolution, the flora and fauna had an essentially modern organization, and grasses as the most important terrestrial producers became widespread during the Miocene. The Miocene radiation of the apes ultimately led to the development of our species, Homo sapiens.
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