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AbstractThis page describes the Precambrian, including stratigraphy, paleogeography, and famous lagerstätten, followed by a sketched outline of some of the major evolutionary events.
Keywords: Precambrian, fossil record, evolution
The Precambrian – i.e., all of Earth's history prior to the beginning of the Cambrian Period – comprises three eons. From oldest to youngest, they are the Hadean, the Archean, and the Proterozoic Eon.
Earth was formed approximately 4600 million years ago. Continental crust had begun to form by ~4200 Ma, and life had evolved – perhaps repeatedly – by 2800 Ma or earlier.
The Hadean Eon was the first in Earth history, extending from the first formation of continental crust, which began some time around 4,600 Ma, becoming persistent soon after, to approximately 4,000 Ma (Cohen et al. 2015). The oldest rocks known are those of the Isua Supracrustal Group of southwestern Greenland, surviving from around 3700 Ma but, unfortunately, preserving no fossils. The Isua rocks are strongly metamorphosed; although some sequences have been demonstrated to be of sedimentary origin, and may have once contained fossils, the heat and pressure to which they have been subjected will have erased any traces.
The earliest terrestrial environments were harsh: Levels of atmospheric oxygen around 1% were too low to sustain an ozone layer, without which there would have been little protection from solar radiation, and high levels of atmospheric carbon dioxide and methane would have created a strong greenhouse effect, with global temperatures estimated to have been between 30 and 50°C (after Willis & McElwain 2002, p. 40).
Oceans formed early, between 4,400 and 3,900, from condensation of atmospheric water vapour. Estimates suggest that the earliest oceans were hot (between 80 and 100°C) and acidic (after Willis & McElwain 2002, p. 40).
Origin of the Earth
Hadean Eon (~4600 to 4000 Ma)
In Europe, the Hadean is overlain by the Ishuan, the boundary being dated around 3,750 Ma. The Precambrian Subcommission, however, does not (at least not yet) recognise the Hadean; the oldest unit recognised by the Subcommission is the Archean, with a lower boundary dated around 3,900 Ma. So defined, the Archean is coeval with the Ishuan and also the uppermost part of the Hadean. Thus there is no simple geochronologic sequence which is recognised everywhere for rocks of this age.
Archean Eon (4000 to 2500 Ma)
approx 3,500 Ma ~ formation of persistent crust
Proterozoic Eon (2500 to ~541 Ma)
Upper (Precambrian-Cambrian) Boundary
Since 1947, when H.E. Wheeler initiated debate with the suggestion that the Precambrian-Cambrian boundary should be based upon the first appearance of trilobites, much has ensued. Progress has largely been facilitated by the International Geological Congress (IGC) and the establishment in 1960 of a Subcommission on Cambrian Stratigraphy. The classical idea of placing the boundary at an unconformity has been displaced by the search for monofacial, continuous deposition sequences across the boundary, with the view to selecting a stratotype.
The search itself produced a wealth of data from around the world – including the Palaeotethyan Belt, Siberian Platform, and England – eventually focusing upon south-east Newfoundland.
Prior to 1990, the boundary was generally placed at the base of the Tommotian Stage, a rock unit known from the Siberian Platform. However, in 1991 the International Subcommission on Cambrian Stratigraphy made a decision to draw the base on the Cambrian at the first appearance datum (FAD) of the distinctive horizontal burrow ichnofossil, Treptichnus (formerly Phycodes and Trichophycus) pedum (Seilacher 1955; Fig. 1), in the reference section at Fortune Head. This horizon correlates with the base of the Siberian Nemakit-Daldynian Stage, ~13 Ma earlier than the Tommotian (included within the Vendian System in some older literature).
Both of the Siberian stages have since been abandoned by the ISCS, however. Presently, the early Cambrian is represented by the Terreneuvian Series, which is divided into the Fortunian Stage (basal) and “Stage 2”. This is unlikely to be the final word on the matter.
Fig. 1: The horizontal burrow trace fossil, Treptichnus (formerly Phycodes and Trichophycus) pedum defines the lower boundary of the Cambrian in the reference section at Fortune Head, southeastern Newfoundland. [Image courtesy of Dr. Gerd Geyer, Institut für Paläontologie, Bayerische Julius-Maximilians-Universität, Würzburg, Germany.]
Major Tectonic EventsThe protracted acceptance of continental drift was one of the great scientific sagas of the twentieth century. Today, although many details of the causal machanisms remain obscure, it is almost universally accepted that the continents ride across the face of the earth on tectonic plates which have been in constant movement since the formation of the crust. The continents we recognise today have not always existed: great blocks of crust have united in many configurations over the planet’s long history. And, more than once, all (or at least the great majority) of the earth’s crust has been united into a single supercontinent.
The most recent such amalgamation was Pangaea, which is relatively well known and about which the interested reader will easily find much information elsewhere. An earlier massing into a single supercontinent during the Precambrian, “remains shrouded in mystery” (Torsvik 2003, p. 1379).
The existence of a Precambrian supercontinent was proposed in the 1970s, largely motivated by the evidence of contemporaneous (1.3 to 1.0 Ga) “Grenville” mountain belts on what are today different continents, and broadly supported by paleomagnetic studies. Various names were offered, including Ur-Pangea, Proto-Pangea and variations (McMenamin 1998, p. 176), but most authorities had already settled on Rodinia by the early 1990s (e.g. Dalziel 1991; Moores 1991; Hoffman 1991).
Note to self: also check Rogers 1996 and McMenamin & McMenamin 1990, p. 95 (MM claims the first usage of the Rodinia name in that pub: McMenamin 1998, p. 176).
Most of the early models sought to align the Grenville-aged mountain belts, with Laurentia forming the core of the continent, flanked by Australia and East Antarctica on the west, and Baltica and Amazonia on the east (fig. 1).
Even today, however, “elucidation of its amalgamation, continental makeup, and fragmentation is hampered by the fact that at any given time, the latitudes for only a few continents are known” (Torsvik 2003, pp. 1380-1381).
By about 1900 Ma, most or all of the continental crust had merged to form a single large continent, named Rodinia, which is thought to have straddled the equator.
Rodinia began forming ~1,300 Ma from the amalgamation of three or four pre-existing continents, in an event known as the Grenville Orogeny, consolidating perhaps 1,100 to 1,000 Ma.
The Grenville Province is a terrane of metamorphism and deformation running from Labrador to Texas, once comprising the southern coast of Laurentia. Ar/Ar dating of gabbros and basalts from the Nova Floresta Formation in western Brazil, parts of the Amazon craton which collided with the region that is now Texas during the formation of Rodinia (Llano Uplift), have yielded dates around 1,200 Ma (Tohver et al. 2002).
Early models assumed that Rodinia remained more or less static from its formation until breaking up, perhaps beginning ~700 Ma, but protracted over many millions of years, into three major blocks: West Gondwana, East Gondwana, and Laurasia (e.g. Rogers 1996).
Australia-East Antarctica rifted away from the western margin of Laurentia first, with other continents rifting away subsequently, and in the case of Baltica and Amazonia, perhaps as late as 600 to 550 Ma, opening the Iapetus Ocean between them. Subsequently – perhaps ~540 Ma – West and East Gondwana collided and merged in the mountain-building event known as the Pan-African Orogeny.
Newer models postulate Rodinia disintegrating before 750 million years ago, most likely between 850 and 800 Ma (Torsvik 2003, p. 1380).
Fig. 2: A conventional reconstruction of Rodinia at 750 Ma (left) and an alternative reconstruction (right) in which the positions of Australia-East Antarctica and Congo have been revised (after Torsvik 2003).
“Geological evidence suggests that Earth’s early atmosphere had little free oxygen, but direct evidence of oxygen content during Earth’s history has been unavailable. Two reports take advantage of the better record of sulfur isotope values for much of the Precambrian to infer the abundance of oxygen and the sulfur chemistry of the early Earth (see the Perspective by Wiechert 2002). Isotopic fractionation is normally mass dependent, but exceptions are known, including the mass-independent photochemical fractionation of sulfur that occurs today in the upper atmosphere where ozone is absent. For sulfur, evidence for mass-independent effects that reflect a paucity of atmospheric oxygen has been found in Hadean crustal rocks. Farquhar et al. 2002, p. 2369, report evidence of mass-independent isotope effects in sulfide inclusions in ancient diamonds, which are derived from deep in Earth’s mantle. These data provide direct evidence that these diamonds were sampling material that reflected processes in Earth’s atmosphere and had been subducted into the mantle during the Hadean. Hadean sulfide rocks are much less depleted than are modern sulfides in 34S, an isotope that reflects bacterial processing. In culture experiments, Habicht et al. 2002, p. 2372, show that the isotope data limit Hadean ocean sulfate concentrations to 200 micromolar or less, considerably lower than previously thought. These data also imply that the atmosphere had little free oxygen but an abundance of the greenhouse-gas methane produced through the action of more prevalent methanogens” (overview from Science, v.298, nr.5602).
Sobolev et al. 2019, p. 555, asserts confirmation of “the mantle source of this water by measuring deuterium-to-hydrogen ratios in these melt inclusions and present similar data for 3.3-billion-year-old komatiites from the Barberton greenstone belt. From the hydrogen isotope ratios, we show that the mantle sources of these melts contained excess water, which implies that a deep hydrous mantle reservoir has been present in the Earth’s interior since at least the Palaeoarchaean era (3.6 to 3.2 billion years ago). The reconstructed initial hydrogen isotope composition of komatiites is more depleted in deuterium than surface reservoirs or typical mantle but resembles that of oceanic crust that was initially altered by seawater and then dehydrated during subduction. Together with an excess of chlorine and depletion of lead in the mantle sources of komatiites, these results indicate that seawater-altered lithosphere recycling into the deep mantle, arguably by subduction, started before 3.3 billion years ago.”
Oxygen was produced by photosynthesis by cyanobacteria (”blue-gree algae”) through the Precambrian, but initially it did not accumulate in the atmosphere. Instead, it combined with iron in the Earth’s oceans to form insoluble iron oxides, which precipitated to the ocean floor as characterisitcally layered “banded iron formations.” The formations are most abundant around 2,400 Ma, and become less common after 1,800 Ma (also see Wikipedia page on Banded iron formations).
Rice et al. 2003 proposes “that there were three Neoproterozoic glaciations. The critical event is the second, Marinoan, glaciation, the most widespread and most easily recognised. Marinoan glacial deposits are overlain by a distinctive transgressive, laminated cap-dolostone, which variably contains isopachous cements, accretionary oscillation megaripples, tubestones, and peloids. δ13C through the cap-dolostone is consistent at -2 to -4L’. The cap-dolostone is bounded above by a flooding surface that corresponds to an increase in the fraction of siliciclastic sediments and, commonly, a shift to from dolomite to calcite. In some successions, seafloor barite and aragonite cements occur at this transition. The Marinoan glaciation, unlike the Sturtian glaciation, is also presaged by a gradual and large (up to 15L’) decline in δ13C. Unequivocal Marinoan deposits include the Ghaub (northern Namibia), Elatina (South Australia), and Ice Brook (northwest Canada) formations, all of which are the upper of a pair of diamictites. Applying these unique isotopic and sedimentological boundary conditions as correlation tools, indicates that the diamictite pairs Petrovbreen + Wilsonbreen (NE Svalbard), Ulvesø + Storeelv (E Greenland), and Surprise + Wildrose (Death Valley) are jointly Marinoan in age. These criteria also indicate that the thick Port Askaig (750m) and Smalfjord (420m) diamictites in Scotland and Norway, respectively, are Marinoan. These correlations are important because, in both cases, the Marinoan diamictite is the lower of two glacial horizons. Thus, it is concluded that the upper diamictites in Norway (Mortenses) and Scotland (Loch na Cille) correspond to a third – Varangerian – glaciation. Varangerian glacial deposits are not widespread, but overlie and appear to be related to the largest Neoproterozoic negative δ13C anomaly (-8L’). This shows up globally between Marinoan and Ediacaran-aged strata (e.g. Wonoka Fm. in S. Australia and Huqf Grp. in Oman). The correlations outlined here imply that glacial deposits corresponding to the earliest – Sturtian – glaciation are absent in Norway, Svalbard, E. Greenland, Scotland and Death Valley. However, cap-carbonates to this glaciation can be recognized in many sequences, based on the isotopic and sedimentological characteristics of the Sturtian cap-carbonates in Namibia (Rasthof), NW Canada (Rapitan), and S Australia (Sturt). In all these cap-carbonates, δ13C rises sharply from -4L’ to +5L’ in relatively organic-rich sediments. Probable Sturtian cap-carbonates, without underlying diamictites, include the lower Russøya Member (Svalbard) and the lower Beck Springs Formation (Death Valley).”
“[O]ceanic redox conditions underwent dynamic fluctuations with several oxygenation episodes of an overall anoxic ocean are proposed for the late Neoproterozoic (Wood et al., 2015; Sahoo et al., 2016; Wei et al., 2018; Zhang et al., 2018). In addition to uncertainties around the trajectory of redox evolution, the triggers for atmospheric and oceanic oxygenation are also disputed. … Tectonic controls on Earth-surface oxygenation have also been invoked in previous research (e.g., Kump and Barley, 2007; Campbell and Allen, 2008; Lee et al., 2016; Planavsky, 2018; Li et al., 2018) although tectonic drivers have not been comprehensively demonstrated by geochemical records. In particular, variations in continental denudation as well as global oceanic circulation following Gondwana amalgamation from the middle Ediacaran to early Cambrian are not well constrained” (Wei et al. 2019).
“There are only three places on Earth with sedimentary rocks older than 3300 million years: the greenstone belts of Isua in southwest Greenland, the Barberton area east of South Africa, and the Pilbara area of northwest Australia. The oldest sediments on Earth have been found in Greenland” (Brack 2001).
Note that these places are vast areas, as opposed to single localities. To illustrate, just one of the several possibly fossiliferous rock units in the Pilbara Craton, the Strelley Pool Formation, is exposed in numerous localities and “its depositional area is assumed to be at least 30,000 km2” (Delarue et al. 2020, p. 2). Unfortunately, it is an infuriating characteristic of the literature dealing with these ancient putative fossil occurrences, that most of the papers (including several cited here) are unprofessionally opaque with respect to exact localities and sample registration details, so it is a major undertaking to figure out just how many different fossil occurrences there are, and which studies apply to which.
The greenstones of the Isua Supracrustal Group of south western Greenland, surviving from around 3700 Ma and possibly more than 3,800 Ma (Nutman et al. 2016), are among the oldest rocks known. They are strongly metamorphosed (to amphibolite facies) although some sequences have been demonstrated to be of sedimentary origin (e.g. Bolhar et al. 2005), possibly even soils (Retallack & Noffke 2019). Until recently, no fossils had been reported from the Isua rocks, due to the high grade metamorphism.
In 2016, Nutman et al. reported 1-4 cm-high stromatolites from “the hinge of an anticline cored by 3,709 ± 9-Myr-old andesitic metavolcanic rocks with locally preserved pillow structures and a maximum metamorphic temperature of ~550 C” (p. 535). This report is certain to be contentious, but if it holds up will be the oldest known fossil report known, predating the Pilbara fossils (see below) by more than 200 million years.
This report aside, evidence that life may have existed in Isua times is indirect, based on the carbon isotope signatures recovered from these ancient rocks. “This isotopic evidence stems from the fact that the carbon atom has two stable isotopes, carbon-12 and carbon-13. The 12C/13C ratio in abiotic mineral compounds is 89. In biological syntheses, the processing of carbon [in] CO2 and carbonates gives a preference to the lighter carbon isotope and raises the ratio to about 92. Consequently, the carbon residues of previously living matter may be identified by this enrichment in 12C. A compilation has been made of the carbon isotopic composition of over 1,600 samples of fossil kerogen (a complex organic macromolecule produced from the debris of biological matter) and compared with that from carbonates in the same sedimentary rocks. This showed that biosynthesis by photosynthetic organisms was involved in all the sediments studied. In fact, this enrichment is now taken to be one of the most powerful indications that life on Earth was active nearly 3.9 billion years ago because the sample suite encompasses specimens right across the geological time scale” (Brack 2001; also see Ueno et al. 2006).
Rosing 1999 report a biological carbon isotope signature from ~3,780 Ma (3,779 ± 81 Sm-Nd date) greywackes and slates with well-preserved sedimentary structures from the Garbenschiefer Formation in the Isua belt.
It should, however, be noted that a number of less enthusiastic commentators advise approaching this interpretation with extreme care (e.g. van Zuilen et al. 2002). For example, Mojzsis et al. (1996, p. 55) claimed to have identified biological carbon isotope signatures from >3,800 Ma aged, chemically precipitated sediments, including banded iron formations (BIFs) and chert, on Akilia Island, southwestern Greenland. However, this interpretation was challenged by Fedo & Whitehouse 2002 who regarded the contested unit as a younger hydrothermal vein. The original claim has been vigorously defended and the final conclusion is as yet unresolved.
3.5 Ga to 3.4 Ga
“The early Archean record tells us that life was present at least 3500 Ma ago. Microbial ecosystems were driven by autotrophy, most likely photoautotrophy, and oxygenic cyanobacteria may already have appeared. Heterotrophs included prokaryotes and, possibly, primitive amitochondrial eukaryotes capable of feeding by phagocytosis. Depending on the amount of O2 available, the biota could also have included aerobic prokaryotes and mitochondrion bearing eukaryotic heterotrophs (but perhaps not eukaryotic algae; see ... Knoll & Holland 1995). Although impossible to test empirically, the possibility that early communities included organisms unlike anything represented in the modern biota cannot be excluded. Clearly, early Archean ecosystems remain poorly understood” (Knoll 1996, p. 55).
However, “[t]he early Archean paleontological record is meagre. Virtually all critical data come from two successions, the Warrawoona Group of Western Australia and the Onverwacht and Fig Tree Groups of South Africa. There may even be redundancy in these two successions, in that some geologists believe that they are tectonically separated portions of a single depositional basin. … “[B]oth successions contain carbonaceous microstructures. These structures are uncommon, and their interpretation as microfossils has been challenged repeatedly (Schopf & Walter 1983; Buick 1991)” (Knoll 1996, p. 53).
The oldest “body” fossils known to date, derive from the Apex Chert, a formation of the Pilbara Supergroup occuring in northwestern Western Australia, and dated at 3,465 Ma (± 5 Ma, see Schopf 1999, p. 88-89). However, the fossils occur in fragments of rock within the chert; thus they are even older, though by how much, is unknown. The organisms themselves are filamentous, composed of distinct, organic walled cells occurring as a uniserial string, and are interpreted as cyanobacteria (Fig. 3). (Schopf 1993, 1999, Altermann & Kazmierczak 2003, Schopf et al. 2018; but see Garcia-Ruiz et al. 2003 for cautions.)
Also occurring in the Pilbara Craton of western Australia, the Strelley Pool Chert (SPC) includes structures interpreted as stromatolites. A 2006 study of these concludes that the evidence “strongly indicates that organisms flourished on a broad peritidal platform 3.43 Gyr ago in the Pilbara, rapidly taking hold and creating a reef-like build-up in shallow waters as surfaces became submerged. The variety of stromatolites present indicates that the SPC may contain not only some of Earth’s earliest fossils but also a diverse fossil ‘ecosystem’, sustained by shallow seawater free of terrigenous influx – ideal conditions for phototrophism” (Allwood et al. 2006, p. 717-718).
“Perhaps a nearly thirty year tradition of rejecting previously reported material while presenting new ‘unequivocal’ evidence is at an end. Schopf (1992c, 1993) has discovered poorly preserved but convincingly biological filaments in cross-bedded Warrawoona chert grainstones. Having visited the outcrop in question, I regard the early Archean age of these fossils as beyond question” (Knoll 1996).
“During the 1970s, further research on Onverwacht and Fig Tree cherts produced a second round of paleontological reports (Muir & Grant 1976; Knoll & Barghoorn 1977). The case for the biogenicity of at least some of these structures is stronger. For example, the structures reported by Knoll & Barghoorn (1977) have a well-defined unimodal size frequency distribution with a mean of 2.5 mm; about 25% of the individuals in a large sample population are clearly paired or have a distinct hour-glass morphology comparable to those of cells undergoing binary fission; the cells have a distinct wall layer and collapsed internal contents, much like that seen in younger microfossils; and individual microstructures may be flattened parallel to bedding, indicating their emplacement prior to sediment compaction. Are they fossils? Quite possibly, but given their simple morphology, unequivocal acceptance of biogenicity is impossible” (Knoll 1996, p. 53).
Although the fossil record for stromatolites stretches back, potentially as far as 3.7 Ga (Nutman et al. 2016), widespread, thick, stromatolitic carbonate platforms appear more or less abruptly between about 2,900 and 2,600 Ma. From one such example, the South African Campbellrand platform, comes another ancient cyanobacterial collection, reported from the Nauga Formation, Prieska, South Africa, and dated between 2,588 ± 6 and 2,549 ± 7 Ma (Kazmierczak & Altermann 2002). From then on, stromatolites provide a virtually unbroken fossil record into the Phanerozoic, including such famous occurrences as the 2,100 Ma Gunflint Formation from Canada and the 850 Ma Bitter Springs Formation in Australia.
“Schopf & Walter (1983) reported rare trichomes from the ca. 2800 Ma Fortescue Group, Western Australia. The fossils resemble oscillatorian cyanobacteria, but they are not taxonomically diagnostic; similar morphologies occur among both sulfur-oxidizing and sulfate-reducing bacteria. More diverse microfossils have been reported from the ca 2500 Ma Transvaal Supergroup, South Africa. Silicified microstromatolites and associated intraclasts from platform environments contain 1-5 mm diameter coccoids and thin filamentous sheaths interpreted as primary producers as well as tiny rods interpreted as heterotrophic bacteria (Lanier 1986). Chert nodules in deeper basinal limestones contain carbonate-lined filamentous sheaths up to 27 μm in cross-sectional diameter (Klein et al. 1987)” (Knoll 1996, p. 55).
Another Late Archean cyanobacteria collection is reported from the Nauga Formation, Prieska, South Africa, between 2588 ± 6 and 2549 ± 7 Ma (Kazmierczak & Altermann 2002).
“Stromatolites shed some light on the nature of early Archean life, but again there are uncertainties and differing interpretations of reported structures” (Knoll 1996).
“[T]he microbial mat origin of the stratiform mats remains well supported. Complex communities of microorganisms including phototactic mat builders, certainly colonized early Archean coastal environments. It is reasonable to suggest that these communities included photoautotrophs, but this is not beyond question (Walter 1983)” (Knoll 1996).
“Stromatolites become increasingly abundant in younger Archean successions, a pattern as likely to reflect craton growth as evolutionary change. By the end of the eon, extensive carbonate platforms supported widespread mat-building communities that almost certainly included cyanobacteria” (Knoll 1996).
The occurrence of diverse [CONFLICTS WITH “HANDFUL OF MORPHOTYPES” BELOW] organisms of this age is consistent with the belief that aerobic metabolism evolved independently in each of the archaeans, true bacteria and eukaryotes, after they had separated from one another, yet prior to the widespread introduction of molecular oxygen into the anaerobic biosphere at approximately 2.2 Ga. “The fact that enzyme distribution in aerobic pathways was largely incongruent with organismal speciation suggests that adaptation to molecular oxygen occurred after the major prokaryotic divergences on the tree of life. This is supported by data from geological and molecular evolutionary analyses, showing that all three domains of life, and many phyla within these domains, had appeared by the time that oxygen became widely available.... The relatively late onset of atmospheric oxidation argues against the invention of O2-dependent enzymes or pathways in the last common ancestor of modern organisms, suggesting that adaptation to molecular oxygen took place independently in organisms from diverse lineages exposed to O2” (Raymond & Segrè 2006).
“The early Archean fossil record at best contains a handful of morphotypes [CONFLICTS WITH DIVERSE ORGANISMA ABOVE], none of which is taxonomically diagnostic or physiologically informative. Indeed the low apparent diversity of early Archean fossils cannot itself be taken at face value. Studies of Proterozoic formations show that with increasing diagenetic/metamorphic alteration, the apparent diversity of microfossil assemblages decreases (Knoll et al. 1988). Thus, when subjected to lower greenschist facies metamorphism, an assemblage with an original diversity comparable to, say, the Gunflint Formation might well yield a morphological record much like that actually seen in early Archean cherts” (Knoll 1996).
Fig. 3: Copy of figures 3I-O from Schopf 1993, showing putative microfossils from the Apex Chert, Western Australia.
At 2,000 Ma we find possible eukaryote fossils, and by 1,750 Ma we are confident of this identification. “An increasingly well resolved Proterozoic fossil record documents the late Mesoproterozoic to early Neoproterozoic presence of most of the major clades (kingdoms) of eukaryotes, including the rhodophytes, stramenopiles, alveolates and green plants. A coincident rise in acritarch diversity, combined with molecular phylogenetic evidence for rapid cladogenesis, points to a major radiation of eukaryote groups at this time, sometimes referred to as the ‘big bang’ of eukaryotic evolution. Bangiomorpha pubescens, from the 1200 Ma Hunting Formation, arctic Canada, is the earliest taxonomically resolved eukaryote on record. More importantly, it is the earliest known example of both sexual reproduction and complex multicellularity. The introduction of differentiated multicellular organisms would have had profound implications for contemporaneous ecology (e.g. with its differentiated basal holdfast Bangiomorpha was able to anchor itself in the substrate and orient itself vertically – the first instance of eukaryotic tiering – which would in turn have induced new environments and evolutionary opportunities). Buss (1987) has presented compelling arguments for why sex is a necessary prerequisite to complex multicellularity (by enabling the expulsion of somatic cell parasites), and it is clear that complex multicellularity is the source of almost all organismal morphology. I suggest that the principal significance of the evolution of sex was the ‘invention’ of organismal morphology, and thereby the directional, escalatory and ‘progressive’ evolution of a biological environment. The modern eukaryotic kingdoms would appear to be the consequence of that first indulgence” (Butterfield 1999a).
By 1,100 Ma we find a rapid diversification of planktonic eukaryote assemblages, followed by an inexplicable decline in both abundance and diversity, between 900 and 675 Ma.
The fact remains, though, that fossils are rare and hard to interpret. Evidence for aggregate biological activity has been inferred from carbon isotope data. “The coevolution of the biosphere and geosphere is reflected in large part by changes in the long term carbon cycle.... Past changes within the cycle are recorded in the isotopic content of carbonate and organic carbon buried in ancient sediments.... Extraordinarily large fluctuations occur in the Neoproterozoic (1,000 to 543 million years ago...) carbon-isotopic record both immediately preceding the Cambrian diversification of complex animal life ... and in the ~200 million years before it.... There is much interest in determining not only the cause of these isotopic events ... but how, if at all, they are related to early animal evolution...” (Rothman et al. 2003, p. 8124).
LagerstättenLagerstätten (sing. lagerstätte) are fossil localities which are highly remarkable for for either their diversity or quality of preservation; sometimes both. Both these criteria are relative and can only be appreciated in some sort of context. In fact, any form of fossil preservation is remarkable in rocks as ancient as these. Thus, although the term is hardly ever applied, there is a case to be made for calling the following fossil beds ‘lagerstätten.’
Apex Chert (Pilbara Supergroup): The oldest plausible fossils reported to date derive from the Apex Chert, a formation of the Pilbara Supergroup occuring in northwestern Western Australia, and dated approximately 3,465 Ma. However, the putative fossils occur in fragments of rock within the chert; thus they are even older, though by how much, is unknown. The structures themselves are filamentous, composed of distinct units occurring as a uniserial string. Some authors interpret these structures to be cyanobacteria; others consider them to be secondary artefacts produced by a hydrothermal system.
Barberton: (= Fig Tree?) Bacteria microfossils dating back 3.3 to 3.4 billion years have also been discovered in rocks from the Barberton greenstone belt, South Africa.
Strelley Group: Long, fine filaments probably representing thermophilic microorganisms living in the vicinity of a hydrothermal vent have been found in a massive sulfide deposit from the Early Hadean Strelley Group (about 3.235 billion years old) of the Pilbara greenstone belt, northwest Australia. Although the temperature of the hydrothermal fluids was about 300°C, the microorganisms more likely developed at temperatures below 110°C and at water depths of about 1000 m. Under such environmental conditions, the microorganisms would have been anaerobic chemotrophs metabolizing in a reducing environment and obtaining their energy and nutrients from the hydrothermal fluids. This deep environment would have provided the microbiota with protection from the harmful UV radiation prevalent at the surface of the Earth during the Hadean, when there was no protective ozone layer (Brack 2001).
The famous ‘Ediacaran’ forms from the latest Precambrian are described in separate pages.
New Zealand Occurrences
Precambrian rocks are almost unknown in New Zealand. Possibly the most likely contender is parts of the Balloon Formation, which is exposed in the Cobb Valley area: see Cooper & Grindley 1982, p. 50.
Allwood, A.C.; Walter, M.R.; Kamber, B.S.; Marshall, C.P.; Burch, I.W. 2006: Stromatolite reef from the Early Archaean era of Australia. Nature 441: 714-718.
Altermann, W.; Kazmierczak, J. 2003: Archean microfossils: a reappraisal of early life on Earth. Research in Microbiology 154: 611-617.
Bolhar, R.; Kamber, B.S.; Moorbath, S.; Whitehouse, M.J.; Collerson, K.D. 2005: Chemical characterization of earth’s most ancient clastic metasediments from the Isua Greenstone Belt, southern West Greenland. Geochimica et Cosmochimica Acta 69 (6): 1555-1573.
Brack, A. 2001: Origin of life. Nature Encyclopedia of Life Sciences.
Brasier, M. 2010: Darwin’s lost world: the hidden history of animal life. Oxford: 1-304.
Buick, R. 1991: Microfossil recognition in Archean rocks: an appraisal of spheroids and filaments from a 3500 m.y. old chert-barite unit at North Pole, Western Australia. Palaios 5: 441-459, pl. 1-6.
Buss, L. W. 1987: The evolution of individuality. Princeton University Press.
Butterfield, N.J. 1999a: Sex, Multicellularity and First ‘Big Bang’ of Eukaryotic Evolution. Palaeontological Association 43rd Annual Meeting, University of Manchester, 19-22 December 1999 (Oral Presentation).
Cohen, K.M.; Finney, S.C.; Gibbard, P.L.; Fan, J.X. 2015: The ICS international chronostratigraphic chart v 2015/01. Episodes 36: 199-204.
Cooper, R.A.; Grindley, G.W. 1982: Late Proterozoic to Devonian sequences of southeastern Australia, Antarctica and New Zealand and their correlation. Geological Society of Australia Special Publication 9: 1-103.
Dalziel, I.W.D. 1991: Pacific margins of Laurentia and East Antarctica-Australia as a conjugate rift pair: Evidence and implications for an Eocambrian supercontinent. Geology 19: 598-601.
Delarue, F.; Robert, F.; Derenne, S.; Tartèse, R.; Jauvon, C.; Bernard, S.; Pont, S.; Gonzalez-Cano, A.; Duhamel, R.; Sugitani, K. 2020: Out of rock: A new look at the morphological and geochemical preservation of microfossils from the 3.46 Gyr-old Strelley Pool Formation. Precambrian Research 336 (105472): 1-12.
Farquhar, J.; Bao, H.; Thiemens, M. 2002: Atmospheric influence of Earth's earliest sulfur cycle. Science 289: 756-758.
Fedo, C.M.; Whitehouse, M.J. 2002: Metasomatic origin of quartz-pyroxene rock, Akilia, Greenland, and implications for Earth's earliest life. Science 296: 1448-1452.
García-Ruiz, J.M.; Hyde, S.T.; Carnerup, A.M.; Christy, A.G.; Van Kranendonk, M.J.; Welham, N.J. 2003: Self-assembled silica-carbonate structures and detection of ancient microfossils. Science 302: 1194-1197.
Habicht, K.S.; Gade, M.; Thamdrup, B.; Berg, P.; Canfield, D.E. 2002: Calibration of sulphate levels in the Archean ocean. Science 298 (5602): 2372-2374.
Hoffman, P.F. 1991: Did the breakout of Laurentia turn Gondwanaland inside-out? Science 252: 1409-1412.
Kazmierczak, J. Altermann, W. 2002: Neoarchean biomineralization by benthic cyanobacteria. Science 298: 2351.
Klein, C.; Beukes, N.J.; Schopf, J.W. 1987: Filamentous microfossils in the early Proterozoic Transvaal supergroup: their morphology, significance, and paleoenvironmental setting. Precambrian Research 36: 81-94, pl. 1.
Knoll, A.H. 1996: Chapter 4. Archean and Proterozoic Paleontology. In Jansonius, J.; McGregor, D.C. (eds.) 1996: Paleontology: Principles and Applications. American Association of Stratigraphic Palynologists Foundation, v. 1, pp. 51-80. 1: 51-80.
— 2003: Life on a young planet. Princeton: 1-277.
Knoll, A.H.; Barghoorn, E.S. 1977: Archean microfossils showing cell division from the Swaziland System of South Africa. Science 198 (4315): 396-398.
Knoll, A.H.; Holland, H.D. 1995: Proterozoic oxygen and evolution: an update. In Stanley, S. (ed.) 1995: Biological responses to past environmental changes. National Academy Press, Washington : 21-33.
Knoll, A.H.; Strother, P.K.; Rossi, S. 1988: Distribution and diagenesis of microfossils from the Lower Proterozoic Duck Creek Dolomite, Western Australia. Precambrian Research 38: 257-279, pl. 1-10.
Lanier, W.P. 1986: Approximate growth rates of early Proterozoic microstromatolites as deduced by biomass productivity. Palaios 1: 525-542, pl. 1-7.
McMenamin, M.A.S. 1998: The garden of Ediacara. Columbia University Press: 1-295.
McMenamin, M.A.S.; McMenamin, D.L. 1990: The emergence of animals. Columbia University Press: 1-217.
Mojzsis, S.J.; Arrhenius, G.; McKeegan, K.D.; Harrison, T,M.; Nutman, A.P.; Friend, C.R.L. 1996: Evidence for life on Earth before 3,800 million years ago. Nature 384: 55-59.
Moores, E.M. 1991: Southwest U.S.-East Antarctic (SWEAT) connection: A hypothesis. Geology 19: 425-428. Geology.
Muir, M.J.; Grant, P.R. 1976: Micorpaleontological evidence from the Onverwacht Group, South Africa. In Windley, B.F. (ed.) 1976: The early history of the Earth. Wiley : 595-604.
Nutman, A.P.; Bennett, V.C.; Friend, C.R.L.; Kranendonk, M.J.V.; Chivas, A.R. 2016: Rapid emergence of life shown by discovery of 3,700-million-year-old microbial structures. Nature 537: 535-538.
Raymond, J.; Segrè, D. 2006: The effect of oxygen on biochemical networks and the evolution of complex life. Science 311: 1764-1767.
Retallack, G.J.; Noffke, N. 2019: Are there ancient soils in the 3.7 Ga Isua Greenstone Belt, Greenland? Palaeogeography, Palaeoclimatology, Palaeoecology 514: 18-30.
Rice, A.H.N.; Halverson, G.P.; Hoffman, P.F. 2003: Three for the Neoproterozoic: Sturtian, Marinoan and Varangerian glaciations. EGS - AGU - EUG Joint Assembly, Abstracts from the meeting held in Nice, France, 6 - 11 April 2003, abstract #11425.
Rogers, J.J.W. 1996: A history of the continents in the past three billion years. Journal of Geology 104: 91-107.
Rosing, M.T. 1999: 13C-Depleted Carbon Microparticles in >3700-Ma Sea-Floor Sedimentary Rocks from West Greenland. Science 283: 674-676.
Rothman, D.H.; Hayes, J.M.; Summons, R.E. 2003: Dynamics of the Neoproterozoic carbon cycle. Proceedings of the National Academy of Sciences of the USA 100: 8124-8129.
Schopf, J.W. 1992c: Paleobiology of the Archean. In Schopf, J.W.; Klein, C. (ed.) 1992: The Proterozoic biosphere: a multidisciplinary study. Cambridge University Press : 24-39.
— 1993: Microfossils of the early Archean Apex Chert: new evidence of the antiquity of life. Science 260: 640-646.
— 1999: Cradle of life: the discovery of Earth’s earliest fossils. Princeton: 1-367.
Schopf, J.W.; Kitajima, K.; Spicuzza, M.J.; Kudryavtsev, A.B.; Valley, J.W. 2018: SIMS analyses of the oldest known assemblage of microfossils document their taxon-correlated carbon isotope compositions. Proceedings of the National Academy of Sciences of the USA 115 (1): 53-58.
Schopf, J.W.; Klein, C. (ed.) 1992: The Proterozoic biosphere: a multidisciplinary study. Cambridge University Press.
Schopf, J.W.; Walter, M.R. 1983: Archean microfossils, new evidence of ancient microbes. In Schopf, J.W. (ed.) 1983: Earth‘s earliest biosphere: its origin and evolution. Princeton University Press, Princeton : 214-239.
Seilacher, A. 1955: Spuren und fazies im Unterkambrium. In Schindewolf, O.; Seilacher, A. (ed.) 1955: Beiträge zur Kenntnis des Kambriums in der Salt Range (Pakistan). Akademie der Wissenschaften un der Literatur, Mainz, Abhandlungen der mathematisch-naturwissenschaftlichen Klasse 10 10: 261-446.
Sobolev, A.V.; Asafov, E.V.; Gurenko, A.A.; Arndt, N.T.; Batanova, V.G.; Portnyagin, M.V.; Garbe-Schönberg, D.; Wilson, A.H.; Byerly, G.R. 2019: Deep hydrous mantle reservoir provides evidence for crustal recycling before 3.3 billion years ago. Nature 571: 555-559.
Tohver, Eric; van der Pluijm, B.A.; Van der Voo, R.; Rizzotto, G.; Scandolara, J.E. 2002: Paleogeography of the Amazon craton at 1.2 Ga: early Grenvillian collision with the Llano segment of Laurentia. Earth and Planetary Science Letters 199: 185-200. earth pl sci lett.
Torsvik, T.H. 2003: The Rodinia jigsaw puzzle. Science 300: 1379-1381. Science.
Ueno, Y.; Yamada, K.; Yoshida, N.; Maruyama, S.; Isozaki, Y. 2006: Ueno, Y.; Yamada, K.; Yoshida, N.; Maruyama, S.; Isozaki, Y. 2006: Evidence from fluid inclusions for microbial methanogenesis in the early Archaean era. Nature 440: 516-519. Nature.
van Zuilen, M.A.; Lepland, A.; Arrhenius, G. 2002: Reassessing the evidence for the earliest traces of life. Nature 418: 627-630.
Walter, M.R. 1983: Archean stromatolites evidence of Earth‘s earliest benthos. In Schopf, J.W. (ed.) 1983: Earth‘s earliest biosphere: its origin and evolution. Princeton University Press, Princeton : 187-213.
Wei, G.; Ling, H.; Shields, G.A.; Chen, T.; Lechte, M.; Chen, X.; Qiu, C.; Lei, H.; Zhu, M. 2019: Long-term evolution of terrestrial inputs from the Ediacaran to early Cambrian: Clues from Nd isotopes in shallow-marine carbonates, South China. Palaeogeography, Palaeoclimatology, Palaeoecology 535: 109367.
Wheeler, H.E. 1947: Base of the Cambrian System. Journal of Geology 55: 153-159.
Wiechert, U.H. 2002: Earth’s early atmosphere. Science 298 (5602): 2341-2342.
Willis, K.J.; McElwain, J.C. 2002: The evolution of plants. Oxford: 1-378.
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