Craig Et Al. 2012 Precambrian Source Rocks JMPG

7/26/2019 Craig Et Al. 2012 Precambrian Source Rocks JMPG

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Review article

The palaeobiology and geochemistry of Precambrian hydrocarbon source rocks

J. Craig a,*,U. Bif a, R.F. Galimberti a, K.A.R. Ghori b, J.D. Gorter a, N. Hakhoo c, D.P. Le Heron d, J. Thurow e,M. Vecoli f

a Eni Exploration & Production Division, Via Emilia 1, San Donato Milanese, 20097 Milano, Italyb Geological Survey of Western Australia, Perth, Western Australia, Australiac Institute of Energy Research and Training, University of Jammu, Indiad Department of Earth Science, Royal Holloway, University of London, United Kingdome Maghreb Petroleum Research Group, Department of Earth Sciences, University College London, United KingdomfUniversit Lille 1, CNRS Research Unit FRE 3298, Gosystmes, SN5 Cit Scientique, 59655 Villeneuve dAscq, France

a r t i c l e i n f o

Article history:

Received 8 May 2012

Received in revised form

21 September 2012

Accepted 27 September 2012

Available online xxx

Keywords:

Precambrian

Hydrocarbons

Source rocks

Palaeobiology

Geochemistry

Petroleum systems

a b s t r a c t

Organic carbon productivity and formation of hydrocarbon source rocks during the Early Precambrian

was almost exclusively the product of the growth of microbial mats. Indirect evidence of microbial mats

can be traced back to at least 2.6e2.7 Ga (Neoarchaean), with the earliest evidence of mat development

in siliciclastic sediments coming from the 2.9 Ga (Mesoarchaean), predominantly marine sedimentary

rocks of the Mozaan Group in South Africa. The earliest direct evidence for terrestrial microbial mats in

siliciclastic sediments comes from the 2.75 Ga (Palaeoproterozoic) uviolacustrine sediments of the

Hardey Formation of the Pilbara craton in Western Australia. Well-preserved Proterozoic hydrocarbons

provide valuable information about the early evolution of the biosphere. Eukaryotic steranes (biomarker

for eukaryotic cells and, therefore, evolved forms of life) are present in the geological record from about

2.7 Ga, but they are not abundant or diverse within Archaean communities, which tend to be dominated

by Archaea isoprenoids. Some hydrocarbons have been generated and migrated from Archaean organic-

rich shales, but they were probably not of sufcient volume to be of commercial interest. The world s

oldest signicant hydrocarbon source rocks are Palaeoproterozoic in age and include the shungite

deposits (2.0 Ga) in the Lake Onega region of Arctic Russia.

There is strong evidence of a global biospheric oxygenation event at c. 1300e1250 Ma (Mesoproter-

ozoic) in conjunction with a rst-order positive shift in the marine carbon isotope record. This is sup-

ported by the appearance of the oldest bedded marine gypsum deposits and of the earliest,

unambiguously multicellular eukaryotes at this time. This oxygenation event probably played a signi-

cant role in supporting the more diverse eukaryotic communities preserved in the Neoproterozoic

molecular record and provided the volume of organic material required to generate commercial volumes

of hydrocarbons. Hydrocarbon source rocks of late Mesoproterozoic and Early Neoproterozoic age are

widespread and include highly organic-rich shales deposited in restricted basinal settings adjacent to

stromatolitic carbonate banks. By c. 850 Ma, the Neoproterozoic molecular record is dominated by

hopanes from cyanobacteria with a signicant abundance and diversity of eukaryotic steranes, including

those of multicellular eukaryotes (red and green algae), as well as molecular evidence for heterotrophic

protists. The most widespread hydrocarbon source rocks of mid to late Neoproterozoic age are commonly

transgressive organic-rich black shales associated with inter-glacial and post-glacial phases of the

Neoproterozoic global scale glaciations. The relative dominance of microbial mats in the contribution oforganic matter as a source for hydrocarbon generation probably decreased signicantly during the late

Neoproterozoic and earliest Cambrian, perhaps as a result of the rapid growth of grazing metazoan

communities or possibly as a result of changes in seawater chemistry and/or nutrient supply.

Precambrian and Infracambrian petroleum systems are relatively abundant and widespread. The

oldest live oil recovered to date is sourced from Mesoproterozoic rocks within the Velkerri Formation

(Roper Group) of the McArthur Basin of northern Australia, dated at 1361 21 Ma and 1417 29 Ma

(ReeOs dates) with at least the initial phase of oil generation and migration having taken place before

1280 Ma. However, the geologically oldest commercial production is currently from the somewhat

* Corresponding author. Tel.: 39 02 520 63596.

E-mail address: [email protected](J. Craig).

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Marine and Petroleum Geology

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http://dx.doi.org/10.1016/j.marpetgeo.2012.09.011

Marine and Petroleum Geology xxx (2012) 1e47

Please cite this article in press as: Craig, J., et al., The palaeobiology and geochemistry of Precambrian hydrocarbon source rocks, Marine andPetroleum Geology (2012), http://dx.doi.org/10.1016/j.marpetgeo.2012.09.011
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younger mid to Late Neoproterozoic (CryogenianeEdiacaran) petroleum systems of the Lena-Tunguska

province in East Siberia and in southern China, from the latest Neoproterozoic to Early Cambrian Huqf

Supergroup in Oman and, potentially in the near future, from the age-equivalent Mawar Supergroup in

western India.

2012 Elsevier Ltd. All rights reserved.

1. Introduction

Interest in Precambrian petroleum systems has gathered pace

over the past decade as they have become increasingly recognised

as a potentially large and relatively untapped resource. The vast

potential of these systems has been demonstrated by the discovery

and exploitation of giant oil and gas elds in China, Russia and

Oman. This has resulted in increasing interest in exploration for

both conventional and unconventional hydrocarbon resources in

Precambrian successions in many areas of the world, particularly in

West Africa, Brazil, North America and Australia (Craig et al. 2009;

Bhat et al., 2012).

True Precambrian Petroleum Systems (as opposed to hybridsystems) require the existence of source rocks of Precambrian age

that are of sufcient quality and volume to generate hydrocarbons

that,through the processesof thermalmaturationand migration, are

subsequently trapped in reservoirs of Precambrian age by seals that

are either of Precambrian or younger in age. In hybridPrecambrian

Petroleum systems either the source or reservoir is generally

younger than Precambrian in age. Most commonly these involve

hydrocarbons generated from Precambrian source rocks that are

trapped in Phanerozoic reservoirs or, alternatively Precambrian

reservoirs that are charged by Phanerozoic source rocks.

Precambrian hydrocarbon source rocks differ from most

conventionalPhanerozoic source rocks in that the organic matter

they contain is predominantly, if not exclusively, of bacterial or algal

origin. Exploration for Precambrian Petroleum Plays requires

a thorough understanding of the both the spatial and temporal

distribution of organic-rich horizons within Precambrian succes-

sions and the progressive evolution of the mix of organic compo-nents within these through geological time. It is the purpose of this

paper to provide a comprehensive review of the palaeobiology and

geochemistry of Precambrian source rocks using a combination of

published information and some new research and analysis and to

track the progressive changes in these through Precambrian time.

1.1. Precambrian stratigraphy and time scale

Precambrian is an informal stratigraphic term that encom-

passes all geologic time from the beginning of the Cambrian Period

(542 Ma) back to the early stages in the formation of Earth. It is

preceded by the informal time unit of the Hadean (Ogg et al.

2008). The Precambrian

is generally subdivided into theArchaean (4000e2500 Ma) and Proterozoic (2500e542 Ma) Eons

(e.g. International Commission on Stratigraphy, 2009, Chart),

although there is a proposal to redene the age ranges of these to

4030e2430 Ma and 2430e542 Ma, respectively (Ogg et al., 2008;

Fig. 1). The Archaean Eon is further subdivided into four Era. The

transition to the Proterozoic is considered to be diachronous in all

cratons and the formalization of a Transition Eon (Eoproterozoic)

is under discussion by the Precambrian sub-commission of IUGS

(Fig. 1). For the present, the Proterozoic Eon conventionally begins

at 2500 Ma; a time of major change in the evolution of the crust,

atmosphere and of life on Earth. The Proterozoic Eon is subdivided

into three Era (Palaeoproterozoic, Mesoproterozoic and Neo-

proterozoic) and further subdivided in 10 Periods, the names of

which broadly re

ect large scale tectonic or sedimentary events.

The term

Infracambrian has, historically, been used in the oiland gas industry to dene the Neoproterozoic to earliest Cambrian

time interval (Smith, 2009). It can also be regarded as including the

Vendian and later Riphean stages of the Russian Neoproterozoic

nomenclature. It was originally proposed as Infracambrien by

Menchikoff (1949)in a paper on the stratigraphy of the Western

Sahara. Subsequently, the term Systme Infracambrien was used

by Pruvost (1951) to include Precambrian sediments underlying

known Cambrian rocks and unconformably overlying generally

metamorphosed strata. Infracambrian remains an informal term,

and is not synonymous with Precambrian. Use of the dened

stratigraphic Periods of Tonian, Cryogenian and Ediacaran is

preferred when stratigraphic dating is robust enough to allow this.

Use of the term Infracambrianis now generally discouraged (e.g.

Craig et al., 2009). For simplicity, the chronostratigraphic divisions

ofPrecambrian timeintoEon, Era and Period are used in this paper

without their relevant sufxes.

Precambrian and Infracambrian (NeoproterozoiceEarly

Cambrian) petroleum systems can be classied as either

producing or proven(those that either do, or could soon, produce

commercial volumes of hydrocarbons) or potential (those where

all the elements of a play are known to exist, but where there is, as

yet, no commercial production). They are relatively abundant and

widespread (Craig et al., 2009;Ghori et al., 2009;Bhat et al., 2012).

1.2. The origin of life on Earth

The fact that life was established on the Earth almost as soon as

conditions permitted the development of a liquid water ocean, hasled to the suggestion that life may have begun in the gas and dust

cloud from which the solar system formed (King, 2009). The Earth

was probably heavily bombarded with meteors until at least 3.85e

3.82 Ga and probably until 3.5e3.0 Ga (Johnson and Melosh, 2012;

Garde et al., 2012) while further periodic heavy bombardments

may have occurred until at least 1.7 Ga (Bottke et al., 2012). Manyof

these bolides probably carried complex organic molecules formed

in space (e.g.Cohen, 1996). Apatite grains from the 3.82 billion year

old Isua Formation of Greenland (Brooks, 2011) have 12Ce13Cratios that are consistent with them having been derived from

living matter (Mojzsis et al., 1996;Mojzsis and Harrison, 2000) and

these chemical signatures may be the earliest signs of biological life

on Earth. The chemical origins of life and the process by which the

earliest life-like

molecules were synthesized remain hotlydebated, but one popular theory invokes a transitional phaseinvolving ribonucleic acid (RNA) e a polymer that is simpler than,

but possibly a precursor to deoxyribonucleic acid (DNA), the

primary molecular building block of life. RNA is considered to be

a potential early-life molecule because it can store information in

its molecular structure, replicate, catalyse the necessary chemical

reactions and because at least one of its major building blocks can

be synthesized from simpler molecules under conditions that could

plausibly have existed on the early Earth (Powner et al., 2009;Van

Noorden, 2009). It is generally, but not universally, accepted that

life existed on Earth as early as 3.5 billion years ago, based on

morphological evidence combined with detailed geochemical and

palaeoecological studies (Czaja, 2010 and references therein).

Sequencing of DNA suggests that the earliest organisms on Earth

J. Craig et al. / Marine and Petroleum Geology xxx (2012) 1e472



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Figure 1. Precambrian hydrocarbon source rocks and Precambrian Stratigraphy.A. Location and stratigraphic age of Precambrian hydrocarbon source rocks described in the text.

Archaean: a. Pilbara Craton, Strelley Poll Chert, Mt. Grant area, Tumbania Fm., Apex Basalt succession, Western Australia; b. Swaziland Supergroup and c. Ntombe Fm. Mozaan Gp.,

South Africa. Palaeoproterozoic: a. Upper Zaonezhskaya Fm., Lake Onega, NW Russia; b. Mugford Gp., Northern Labrador; c. Tyler Fm./Michigamme slate, N Michigan; d. Foslev Fm.,

Sortis Gp., SW Greenland; e. Gabon; f. Ontario, Canada; g. Great Lakes Region, Central North America; h. Changcheng Gp., North China; i. Stirling Range Fm., Western Australia; j.

Duck Creek Dolomite, SW Pilbara, Australia. Mesoproterozoic: a. Velkerri Fm. McArthur Basin, Australia; b. Bangemall Gp., NW Australia; c. Siberian Platform, Russia; d. Atar Gp.,

Toudenni Basin, West Africa; e. Sirban Limestone, NW Himalaya, India; f. Vindhyan Basin; g. Chattisgarh Basin; h. Cuddapah Basin. Neoproterozoic: a. Beck Spring Dolomite,

southern California; b. Tindouf Basin, southern Morocco; c. Pertatataka Fm., Amadeus Basin, d. Ungoolya Gp., Ofcer Basin, Australia; e. Doushantuo and Dengying Fms., South

China; f. Vacheda Fm., East European Platform; g. Chuar Gp., Arizona, USA; h. East Svalbard/East Greenland Platform; i. Bitter Springs chert, Central Australia; j. Tindir Gp., Alaska; k.Vazante Gp., SE Brazil; l. Centralian Superbasin, Australia; m. Tapley Hill Fm., S Australia; n. Katanga region, Congo; o. S Oman Salt Basin; p. Gammon Ranges, S Australia; q. Makenzie

Mountains, Canada; r. Nyanga-Niari Basin, S Gabon; s. Marwar Super gp., Rajasthan, India; t. Salt Range Fm., Pakistan; u. Huqf Super Gp., Oman; v. Glacial deposits, Namibia.B. The

current International Stratigraphic Chart for the Precambrian (left) and possible changes to the Precambrian time scale under consideration (right) (after Ogg et al., 2008).

J. Craig et al. / Marine and Petroleum Geology xxx (2012) 1e47 3

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were thermophilic (i.e. they could thrive in relatively high

temperatures between 45 and 80 C), allowing them to survive in

oceans or pools that were heated by volcanic activity, hot springs

and bolide impacts.

It is widely accepted that the isotopic ratio of 12C/13C in

organic carbon is the result of metabolic processes triggering

a fractionation effect on carbon isotopes. Metabolic processes

produce distinctive isotopic fractionations when selecting carbon

from chemical (CO2) or organic substrates. Increased ratios of

12C/13C (expressed as d13C values) can be used as a proxy indi-

cator for early life/photosynthetic processes (e.g. Schidlowski, 1988,

2001). The isotopic signature of organic carbon in sedimentary

rocks is particularly useful to understand variations in microbial

activity (Nisbet and Fowler, 1999; Grassineau et al., 2001, 2002) and

thed13C signature of molecular remains can indicate the presence

of a particular form of life (review inNisbet and Sleep, 2001).

1.3. Precambrian hydrocarbon source rocks

For the purpose of this review, hydrocarbon source rocks are

dened as rocks from which hydrocarbons have been generated or

are capable of being generated. They are organic rich (typically with

more than 1% Total Organic Carbon) and may have been depositedin a wide variety of different deep marine, lacustrine and deltaic

environments. The palaeobiology and geochemistry of these

organic-rich rocks and how these change through the Precambrian,

provides critical information about their environment of deposition

and their suitability to generate hydrocarbons during thermal

maturation.

Biomarker analysis is a key tool in understanding the biological

character of Precambrian source rocks and in determining the

contribution made by different organisms to the organic matter

preserved in them. Organic triterpenoids extracted from source

rocks and oils can be considered as molecular fossils derived from

biological precursor compounds. Steranes are derived from

eukaryotic organisms (i.e. those with a cell nucleus), mainly algae

and higher plants, whereas hopanes (terpanes) are derived fromprokaryotic organisms (i.e. those lacking a cell nucleus) in the form

of bacteria (Rollinson, 2007). Particular carehas to be taken toavoid

contamination during biomarker analysis which can introduce

spurious biomarkers into thermally overmature and/or organic-

lean samples: Contamination with reworked biomarkers is

a particular problem since distinctive biomarkers, even those

associated with the same species, transform through geological

time.

2. Archaean (Pre-2500 Ma): the dawn of life

The search for the earliest forms of life on Earth is difcult,

largely because the few Archaean sedimentary rocks that have

survived are usually severely altered by metamorphism. As a result,the Archaean palaeobiological record is rather meagre. The main

documented Archaean microfossils are carbonaceous bodies,

colonial unicells and lamentous structures. Many of these have

been interpreted as the remains of cyanobacteria (a phylum of

bacteria, formerly called blueegreen algae, that contain blue

pigment in addition to chlorophyll and which obtain their energy

through photosynthesis) and cysts ofagellates.

2.1. The worlds oldest fossils?

The oldest purported fossils are from a c. 3465 million year old

beddedchert unit within the Early Archaean Apex Basalt succession

of the Pilbara Supergroup in northwest Western Australia (Schopf,

1993, 1999;Fig. 2). This apparent prokaryotic assemblage has been

interpreted to include rather advanced forms of cyanobacteria

(Schopf and Kudryavtsev, 2010) of at least eleven different taxa,

preserved as lamentous, dark brown to black carbonaceous

microfossils. The irregularly distributed and randomly oriented

solitary cylindrical laments are surrounded by homogeneous

brown to dark brown kerogen, which may originally have been

mucilaginous. It has been claimed that these Early Archaean rocks

also contain solitary unicell-like spheroids of possible, but uncer-

tain, biological afnity (Schopf, 1993).

The existence of the Apex Chert fossils, if genuine, would imply

that oxygen-producing, photosynthesizing, cyanobacteria devel-

oped rather early in the history of the Earth and even before the

widely accepted great oxygenation event, between 2450 and

2200 Ma,during which the Earths atmosphere changed from being

anoxic to being weakly oxic (Cloud, 1972; Holland, 2002;McCall,

2009b; Gaillard et al., 2011). This observation is important

because the ability of cyanobacteria to perform oxygenic photo-

synthesis is considered to be a key factor in the conversion of the

Earths early reducing atmosphere into an oxidising atmosphere,

thereby provoking an explosion of biodiversity and simultaneously

leading to the near extinction of oxygen-intolerant organisms. The

organic nature of the Apex Chert microfossils has, however, been

strongly disputed. The fossil-like structures have also been inter-preted as artefacts formed from amorphous graphite within

multiple generations of metaliferous hydrothermal vein chert and

volcanic glass (Brasier et al., 2002). Further detailed analysis has

suggested that the Apex Chert laments are possibly fractures

lled with haematite and quartz and not necessarily biological in

origin, but it has also revealed the presence of carbonaceous

material in the matrix of the chert that could be organic, so the

Apex Chert may yet contain evidence of very early life (De Gregorio

et al., 2009;Schopf and Kudryavtsev, 2010;Marshall et al., 2011).

2.2. Other Archaean microfossil assemblages

Other, relatively well established Archaean microfossil assem-

blages include sheath-enclosed colonial unicells in the c. 3465million year old sedimentary rocks of the Towers Formation,

immediately underlying the Apex Chert (Schopf and Parker, 1987),

narrow bacterium-like laments from the c. 3450 million year old

Swaziland Supergroup of South Africa (Schopf, 1992;Walsh, 1992;

Walsh and Lowe, 1985), probable stromatolites in the 3430 million

year old Strelley Pool Chert (Allwood et al., 2006) and a variety of

microfossils from c. 3000 Ma rocks in the Mount Grant area, both

within the Pilbara Craton, Western Australia (Grey and Sugitani,

2009) and two types of cyanobacteria-like laments from the c.

2750 million year old Tumbiana Formation, also in Western

Australia (Schopf and Walter, 1983). The light carbon isotopic

values obtained from the graphitic Apex Chert support the exis-

tence of a signicant biological contribution to the carbon cycle at

this early stage in the Archaean. Given the hydrothermal nature ofthese chert deposits, the most likely source of this biological

component is probably hyperthermophile bacteria, which are

known to occur in younger Archaean rocks (Westall et al., 2001;

Brasier et al., 2002). This would be consistent with the sequencing

of bacterial ribosomal ribonucleic acid (RNA), which suggests that

methanogenic archaeobacteria have a much longer evolutionary

history than cyanobacteria (Hedges et al., 2001; Brasier et al., 2002).

2.3. Biological and chemical evidence for oxygenation of the early

atmosphere

The Marble Bar Chert member of the Duffer Formation, Pilbara,

Western Australia, which underlies the Apex Basalt, appears to

contain primary haematite (i.e. not an oxidation product of




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magnetite or siderite). This would suggest that oxygenated

seawater existed locally, at least at times, during the early Archaean

before the Apex basalt was erupted at c. 3460 Ma (Hickman, 2009;

Hoashi et al., 2009). The isotopic, mineralogical and elemental

evidence from the Marble Bar Chert are all consistent with the

presence of microbes that were capable of producing oxygen (Czaja,

2010). Large (up to 300 mm) spheroidal carbonaceous micro-

structureshave also been described from the 3200 Ma sedimen-

tary rocks of the Moodies Group of South Africa. These have been

tentatively interpreted as organic-walled microfossils of uncertain

biological afliation, which would classify them as acritarchs

(Javaux et al., 2010). If this is conrmed by further studies, it would

imply that large micro-organisms already existed in the photic zoneof marine environments, cohabiting with benthic microbial mats,

during Archaean times. At present, the nature and signicance of

these carbonaceous structures, however, remain enigmatic, and

their palaeoecological interpretation is largely speculative. The

existence of some oxygen in the Earths atmosphere at this early

stage is also hinted at by elemental-speciation and carbon isotope

data. For example, it has been suggested that manganeseereducing

bacteria may have contributed to the formation of 2.92e2.96 Ga

manganese-rich carbonate deposits in South Africa. Manganese-

reducing bacteria require a source of oxidized manganese. The

only known mechanism for producing this requires free oxygen,

but it is unclear whether this implies that there were local oxygen

oases or more global oxygen pulses at this time (Czaja, 2010).

Pervasive oxygenation may have occurred preferentially along

organically-productive Archaean ocean margins such as the

CampbellrandeMalmani carbonate platform in South Africa, where

the 2.67e2.46 Ga succession of the Transvaal Supergroup contains

organic-rich black shale beds up to 20 m thick with Total Organic

Carbon (TOC) contentsof between 2% and 5%, and peaks in excess of

10% (Kendall et al., 2010). These could originally have been some of

the worlds oldest potential hydrocarbon source rocks. Redox-

sensitive metal enrichment (rhenium and molybdenum) together

with the sulphur and nitrogen isotope signatures of these black

shales suggest the presence of dissolved oxygen in the bottom

waters, below the photic zone, with mildly oxygenated surface

ocean waters and an anoxic deep ocean locally at this time (Kendall

et al., op cit.).It is has been proposed that, during the late Archaean, the Sun

was approximately 20% dimmer than at present, but that in

combination with a somewhat denser atmosphere (Som et al.,

2012), containing higher concentrations of greenhouse gases

(Kasting, 1987; Haqq-Misra et al., 2008), it warmed the Earth

sufciently to give a temperate climate (Kasting and Howard, 2006;

Hren et al., 2009) and to allow liquid water to exist at the surface

(the so-called Faint Young Sun paradox). Classical sea-oor

spreading and the development of deep oceans and mid-ocean

ridges only began at the end of the Archaean when the crust had

cooled enough for deep fracturing to allow plates to separate and

move apart (McCall, 2010;Eriksson et al., 2012). The abundance of

Banded Iron Formationsin the Archaean record has been taken as

circumstantial evidence for at least partial oxygenation of the

Figure 2. Putative microfossils from the c. 3.465 Ga Early Archaean Apex Chert, Pilbara Supergroup, northwestern Western Australia, as originally identi ed and illustrated by

Schopf (1993).A. Location Map.B. Stratigraphic column, distribution of reported stromatolites and microfossils, approximate ages and carbon isotope data for geological units of the

Pilbara Supergroup.C. Putative microfossiliferous (1 and 2), laminated stromatolite-like clasts (3) and carbonaceous and ironestained (12) microfossils (with interpretive drawings)

shown in thin sections. Magnication denoted by the scale in (14) unless otherwise indicated. (1) Microfossiliferous clast; area denoted by dashed lines illustrated in 2. (2). Arrows

indicate the positions of minute lamentous microfossils randomly oriented in the clast. (3). Portion of a clast showing stromatolite-like laminae. (4) and (5). Archaeotrichion

septatum, n.sp. (6).Eoleptonema apex, n. sp. (7) and (8). Primaevilum minutum, n. sp. (9), (10) and (11). Primaevilum delicatulum Schopf, 1992. (12), (13), (14) and (15). Archae-

oscillatoriopsis disciformis, n. gen., n. sp.




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oceans, since their development would have required a reliable and

substantial source of oxygen. Whether this oxygenation was the

result of cyanobacterial activity remains to be proven (Morton,

2009). Various other causes for the great oxygenation eventhave been proposed, including changes in the oxygenation state of

volcanic gases as volcanic activity changed from dominantly

submarine to widely subaerial as continental crust began to form at

the end of the Archaean (Gaillard et al., 2011 and references

therein). Certainly, oxygen produced as a waste product of cyano-

bacterial photosynthesis could have combined with ferrous iron

dissolved in marine water to form insoluble ferric oxide which

precipitated as banded iron formations, but many other mecha-

nisms have been suggested for the formation of these deposits,

including photodissociation of water in the atmosphere, photo-

chemical reactions, the photooxidation of ferrous iron and the

actions of an oxygenic phototrophic bacteria (e.g. Cairns-Smith,

1978; McNamara and Awramik, 1994; Widdel et al., 1993). It is

unlikely that appreciable amounts of oxygen would have accumu-

lated in the atmosphere at this time because the rate of photo-

synthetic oxygen production was probably insufcient to overcome

the oxygen consumption as a result of crustal weathering or by

reaction with reduced gases in the atmosphere and the anoxic deep

oceans (Kendall et al., 2010).

2.4. Therst microbial mats

In modern siliciclastic environments, benthic microbiota form

microbial mats that carpet the sea oor. These microbial mats are

composed of bacterial cells and their mucus of extracellular poly-

meric substances. Indirect evidence of ancient microbial mats can

be traced back to at least 2700e2900 Ma (Watanabe et al., 2000).

Organic carbon productivity during the Early Precambrian was

probably almost exclusively the result of the growth of microbial

mats (Schopf, 1999). The earliest evidence of mat development in

siliciclastic sediments comes from the 2900 Ma shallow marine

sandstones of the Ntombe Formation, Mozaan Group in South

Africa (Noffke et al., 2003; Fig. 3). These sandstones preserve

wrinkle structures which, in thin section, exhibit lamentous

textures forming carpet-like microbial mat fabrics that resemble

the laments (trichomes) of modern cyanobacteria (lamentous

cyanobacteria reproduce by fragmentation of their laments),

chloroexi (a bacterial phylum) or sulphur-oxidizing proteobac-

teria. Mineralogical, geochemical and isotopic analyses are consis-

tent with these lament-like textures being of biological origin.

Organic carbon appears to line the former cell walls of the

trichomes (Fig. 3D).The carbonlaments have an isotopic signature

(d13C 24.2 per mil 0.5 per mil) that is consistent with

Figure 3. Microbial mats in the siliciclastic rocks of the c. 2.9 Ga Mesoarchaean Mozaan Group, Pongola Supergroup, South Africa, as illustrated and described by Noffke et al.

(2003). A. Location of the Mozaan Group outcrops in South Africa. B . Location of the sections of the Ntombe Formation. C . Stratigraphic setting of the Ntombe Formation. D .

Elemental composition and abundances of the mat fabrics showing that the inner parts of the laments are composed of iron-oxides (Fe), whereas their outer walls are line with

carbon (C).E. Mat fabrics in the ne sandstones of the Mozaan Group, as seen in thin section (photograph and sketch). The lament-like microstructures (f) resemble trichomes of

bacteria or cyanobacteria. They appear to trap detrital quartz grains (qu) and in situ-formed chert particles (ch) and construct carpet-like fabrics characteristic oflamentous mats.In the sketch, the dark areas (cc) are cuts through laments.F. Wrinkle structures on the bedding surface of a ne sandstone from the Mozaan Group. G. Thin section ofne

sandstone from the Mozaan Group showing dark laminae comprising lamentous microstructures representing microbial mat layers (MM) alternating with sandy layers with

secondary porosity. The pores (P) result from the pressure of gas trapped beneath the sediment-sealing microbial mat layers.




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a biological origin (Schidlowski et al., 1983) and which contrasts

with the isotopically heavier range typical of Archaean inorganic

graphitic material (Noffke et al., 2003). The carbon laments are

closely associated with haematite, goethite and chert which may

derive from the former presence of oxygen within the microbial

mats. Microbial mats that form in siliciclastic depositional envi-

ronments are often subject to erosion by waves or currents and

their fabric must be sufciently coherent to resist mechanical

destruction. The dense microbial mats that form in these envi-

ronments typically cover and stabilise the underlying sediment

(Paterson, 1994). In benthic marine environments, the survival of

light-sensitive microbiota depends on their ability to incorporate

mineral particles into their fabrics by bafing, trapping and

bindingat approximately the same rate as sediment is deposited

(Black, 1993; Noffke, 2009). The microbial mats in the Mozaan

Group contain detrital quartz, zircon and rutile (Fig. 3E), which

would be consistent with the bafing, trapping and binding

process. Together with the hydrodynamic conditions under which

the Mozaan Group sediments were deposited, this suggests that the

biota responsible for the mats were photoautotrophic cyanobac-

teria, although it would not exclude the possibility that they were

sulphur-oxidizing proteobacteria.

Mesoarchaean life is recorded by the body fossils of microbes incherts and stromatolites in carbonate rocks from marine environ-

ments where chemical precipitation led to the rapid lithication of

the bacterial cells. The bacterial groups to which these fossils

should be attributed are, however, largely unknown. In modern

marine environments, benthic bacteria that form laments of

comparable size to those preserved in the 2900 Ma Mozaan Group

sediments include photosynthesising cyanobacteria and sulphur-

oxidizing proteobacteria (Noffke et al., 2003). Stromatolites

become increasingly abundant in the Neoarchaean and it is clear

that extensive carbonate platforms which developed at this time

supported widespread mat-building communities that almost

certainly included cyanobacteria.

The earliest direct evidence for terrestrial microbial mats in

siliciclastic sediments comes from the 2750 Ma uviolacustrinesediments of the HardeyFormation (Fortescue Group) of the Pilbara

Craton in Western Australia, which contain millimetre-sized

pendant columnar structures (with stromatolitic lamination) in

syn-sedimentary cavities (Rasmussen et al., 2009), and from the

1800 Ma palaeo-desert sediments of the Waterberg Group in South

Africa (Eriksson et al., 2000). The syn-sedimentary cavities

preserved in the Hardey Formation resemble the open voids that

develop between dense microbial mats and underlying tidal sands

in modern environments (Noffke et al., 2001). In these modern

environments, the cohesive mats act as impermeable seals to

upward-migrating gas (typically methane) leading to doming of the

mats and the development of sheet-like gas-lled hollows (Gerdes

et al., 2000; Rasmussen et al., 2009) and the syn-sedimentary

cavities in the Hardey Formation may have formed in a similarway. The identity of the microbes responsible for the mat devel-

opment in the Hardey Formation is not known, but their cavity-

dwelling habit suggests that they were not photosynthesising

forms. Carbon and sulphur isotope analyses suggest that the cavity-

lling biolms were probably inhabited by methane and sulphur-

metabolizing bacteria (Rasmussen et al., 2009).

2.5. Geochemical evidence of Archaean life

There is good evidence that life was ourishing in Archaean

oceans by about 2700 Ma (Nisbet and Sleep, 2001;Schopf, 2006)

although body fossils of bacteria, biogenic sedimentary structures

and biomolecules are only rarely preserved (Knoll, 1999; Brasier

et al., 2002;Schopf et al., 2002). Steranes, molecules with 26e

30

carbon atoms arranged in four rings, produced by the decay of

cholesterol and other steroids found in the membranes of

eukaryotic organisms (e.g. algae), are present in the geological

record from about 2700 Ma (Brocks et al., 1999;Summons et al.,

1999), but they are not abundant or diverse in Archaean rocks,

which tend to be dominated by archaeobacterial isoprenoids

(Schuneman et al., 2002; Ventura et al., 2007; Waldbauer et al.,

2009). Bitumen extracts from cores drilled in the 2.67 to 2.46 Ga

Transvaal Supergroup in South Africa contain a suite of molecular

fossils that include hopanes attributable to bacteria, probably cya-

nobacteria and methanotrophs, together with steranes of eukary-

otic origin (Waldbauer et al., 2009) and so support the existence at

this time of both multicellular life as well as of oxygenetic photo-

synthesis and the anabolic use of oxygen. Similar molecular fossils

have also been recovered from the 2.78 to 2.45 Ga old black shales

belonging to the Mount Bruce Supergroup in the Hamersley Basin

of Western Australia (Brocks et al., 2003), but the indigenous origin

of many Archaean hydrocarbon biomarkers and their role as indi-

cators of the presence of early oxygen-producing organisms have

both been challenged (Rasmussen et al., 2008) on the basis that

similar compounds can also be produced by abiotic (i.e. non-

living) processes under hydrothermal conditions (McCollom and

Seewald, 2006) and by anoxygenic photoautotrophs e

bacteriathat carry out photosynthesis to acquire energy and can x carbon,

but without the production of oxygen (Rashby et al., 2007;Kendall

et al., 2010).

2.6. Presence and effectiveness of Archaean hydrocarbon source

rocks

There is convincing evidence that some Archaean rocks were

sufciently organic-rich to have generated and expelled hydrocar-

bons during their subsequent history (e.g.Buick et al., 1998). While

some thick, organic-rich black shale units of Archaean age that

could originally have been potential volumetricallyesignicant

hydrocarbon source rocks do exist, they are rare. There is no

evidence that the quantities of hydrocarbons released fromArchaean source rocks and trapped in reservoirs were sufciently

large and/or were preserved long enough to represent commer-

cially signicant accumulations today although the existence of

some Archaean generated hydrocarbons is indicated by the pres-

ence of rare pyrobitumen nodules of migrated oil and of hydro-

carbon inclusions in some less deformed and metamorphosed

Archaean rocks (Peters et al., 2005).

3. Palaeoproterozoic (2500e1600 Ma): the worlds oldest

petroleum source rocks

3.1. Increased biodiversity and the great oxygenation event

The fossil record is much more complete and relatively contin-uous from about 2100 Ma, beginning with the diverse microbiota

recordedfrom the c. 2100 Ma Belcher Group (Hofmann,1976) and c.

2080 Ma Gunint Iron Formation (Barghoorn and Tyler, 1965) in

Canada. Filamentous sheaths and small simple ellipsoidal fossils

and dyads and quartets of cocci (any spherical or nearly spherical

bacteria) are relatively common in the Palaeoproterozoic and are

widely considered to be the remains of cyanobacteria and eubac-

teria. The recent discovery of large and complex colonial organisms

in 2.1 Gyr black shales in Gabon (El Albani et al., 2010) has changed

the conventional perception of low-diversity prokaryote-domi-

nated palaeoproterozoic marine ecosystems. The richness and

morphological diversity of the Gabon macrofossils clearly indicates

that there was a period of intense taxonomic biodiversication of

marine life during the Palaeoproterozoic, signi

cantly pushing back




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in time the rise of rst representatives of multicellular life. The

existence of such large organisms in Palaeoproterozoic times also

implies that levels of biological productivity in oceanic waters must

have been much higher than previously thought in order to sustain

such complex ecosystems. This 2.1 Ga biodiversication occurred

after the Great Oxygenation Event (2.45e2.32 Ga; Bekker et al.,

2004) during which oxygen levels in the atmosphere are thought

to have increased from below 10-5 times the present atmospheric

level to 10-1 to 10-2 times present day levels (Czaja, 2010).

Although biodiversity increased steadily during the early

Palaeoproterozoic, the total volume of organic matter that could be

preserved seems to have been insufcient to generate viable

hydrocarbon source rocks until c. 2.1e2.0 Ga when the total volume

of organic matter that could be preserved in the geologic record

appears to have increased dramatically. This is entirely consistent

with the c. 2.0 Ga age of oldest known organic-rich hydrocarbon

source rocks deposited in the UpperZaonezhskaya Formation in the

Lake Onega area of northwest Russia (Fig. 4).

3.2. Russian shungite deposits: the worlds oldest signicant

hydrocarbon source rocks?

This Palaeoproterozoic succession contains one of the mostremarkable accumulations of organic carbon in the Precambrian

geological record. The carbon occurs in the form of shungite(named after the small local village of Shunga in the Lake Onega

region), a dense, black, amorphous or non-crystalline, non-

graphitised, structurally-heterogeneous, semi-metallic, glassy

mineraloid that contains >98 wt% of carbon with traces of

nitrogen, oxygen, sulphur and hydrogen (Melezhik et al., 1999;

Fig. 5). The shungite occurrences represent a mixture of meta-

morphosed oil shales, a breached palaeo-oil eld, fossilised

organo-siliceous diapirs and palaeo-oil seeps. The original

organic matter was deposited in an oxygenated, sulphate-poor,

brackish-water lagoonal environment in a volcanic continental

rift setting on the margin of the Archaean Karelian craton

(Melezhik et al., 1999, 2004). In situ, stratied deposits ofshungite represent metamorphosed oil shales (80 wt % C).

The diapiric deposits form non-stratied cupolas or mushroom-

shaped bodies (Fig. 6) composed of impure shungite (20e55 wt

% carbon and 35e75 wt. % silicon dioxide). These are inter-

preted to be organo-siliceous rocks, probably originally gels or

muds. Shungite also occurs as


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Figure 4. Palaeoproterozoic shungite and shungiteebearing rocks from the Lake Onega region, northwest RussiaA. Geological map of the northern Lake Onega region of northwest

Russia showing the location and geological setting of the Palaeoproterozoic shungite deposits (afterMelezhik et al. 2004).B. Lithological subdivisions of the northern Lake Onega

region showing the stratigraphic position of the shungite layers. C. Composite lithological section of the Shungskoe deposit (afterMelezhik et al. 2004).




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The oldest undisputed acritarchs (excluding the debated

occurrence of 3200 Ma acritarchs reported by Javaux et al., 2010

and discussed above) are from shales and silty shales within the

c. 1800 Ma Changzhougou Formation (Lamb et al., 2009) and the

overlying 1700 50 Ma Chuanlinggou Formation of the Chang-

cheng Group (Peng et al., 2009) in North China. In fact, this is the

only known occurrence ofbona de Palaeoproterozoic acritarchs.

Although the Changcheng Group acritarchs are relatively simple

sphaeromorphs, their large size and ne-scale morphological

complexity (presence of medial splitting, variation in cell wallstructure and wall exibility) indicates that they have strong

eukaryotic afnities and suggests that the Eukarya rst diversied

during the Palaeoproterozoic (Lamb et al., 2009). There are also

simple disks and millimetre-scale trails preserved in the tidally-

inuenced marine sandstones from the 2.0 to 1.8 Ga Stirling

Range Formation in Western Australia (Rasmussen et al., 2004)

which suggest that some of the earliest eukaryote animals that

existed at this time may even have been capable of limited move-

ment across the sea oor (Bengtson and Rasmussen, 2009;

Rasmussen, 2009;McCall, 2009a). Most acritarchs are interpreted

as unicelled photosynthetic protists (Martin, 1993;Strother, 1996),

although some may be multicellular algae (Mendelson, 1987;

Buttereld, 2004) and a few have been tentatively interpreted as

fungi (Butter

eld, 2005).

Until the Late Neoproterozoic (Ediacaran), the acritarch record

was dominated by a low diversity assemblage of exceptionally

long-lived taxa, primarily unornamented sphaeromorphs, together

with some acanthomorphic forms with irregularly distributed

processes (Knoll, 1994;Buttereld, 1997,2007;Javaux et al., 2003;

Peterson and Buttereld, 2005;Huntley et al., 2006; Knoll et al.,

2006; Buttereld and Grotzinger, 2012). If most acritarchs are the

cysts of unicellular phytoplankton, they were probably widely

distributed and relatively unconstrained by local environmental

conditions, but there is increasing evidence that Proterozoic acan-thomorphs are the remains of benthic, heterotrophic and/or

multicellular organisms which did occupy a particular ecological

niche, in predominantly shallow water environments (Buttereld

and Grotzinger, 2012). Although such strong ecological partioning

does not preclude the use of acritarchs as biostratigraphic markers,

it does have important implications for the interpretation of

depositional environments and biofacies.

3.5. Presence and effectiveness of palaeoproterozoic hydrocarbon

source rocks

In summary, although there are no present-day commercial oil

and gas accumulations sourced from Palaeoproterozoic sediments,

rocks of this age do contain abundant microfossils, organic matter

Figure 5. Major types of Palaeoproterozoic shungite and shungiteebearing rocks from the Lake Onega region, northwest Russia (as illustrated by Melezhik et al., 2004). A.

Semilustrous shungite rock with well-developed parting from the Shungskoe mine. Width of photograph 1 m. B . Semimat shungite rock with weak parting. Note poorlydeveloped lamination in the upper part where the rock contains higher abundances of siliciclastic material, Shungskoe mine. Width of photograph 1 m.C. Lustrous layer-shungite

(dark brown) sandwiched between a lens of diagenetic dolostone (light grey, above) and semilustrous shungite (beneath). Dolomite concretion is located beneath the layer of

lustrous shungite (beneath 25 cm knife), Shungskoe mine. D. Lustrous layer-shungite made up of 98.4 wt.% Corg with jarosite lms (brown) from the Shungskoe mine. Width of

specimen 5 cm.E. Massive shungite rock from the Maksovo deposit.10 cm scale bar.F. Quartz-cemented shungite rock breccia from the Maksovo deposit. 10 cm scale bar.G. Vein-shungite cross-cutting semimat shungite rocks from the vicinity of the Maksovo deposit. 10 cm scale bar. H. Disc-like inclusion of lustrous shungite (pyrobitumen or anthraxolite)

spread along the bedding surface of thinly laminated Kondopozhskaya siltstones in Nigozero quarry. Shungite discis intensively joined as a result of shrinkage. 10 cm scale bar.I.

Several clasts of lustrous shungite in the matrix-supported conglomerate of the Kondopozhskaya Formation, Nigozero quarry. 10 cm scale bar.




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and remnant hydrocarbons and there is evidence from the Karalian

shungite in Russia that hydrocarbons were generated and migrated

from some organic-rich Palaeoproterozoic rocks, at least locally, in

the past.

4. Mesoproterozoic (1600 Mae1000 Ma): the worlds oldest

commercial oil and gas elds

4.1. The worlds oldest live oil

The oldest live oil recovered in the world to date is sourced from

early Mesoproterozoic rocks within the Velkerri Formation (Roper

Group) of the McArthur Basin of Northern Territory, Australia

(Jackson et al., 1986; Crick et al., 1998; Dutkiewicz et al., 2007;

Fig. 9). The McArthur Basin contains an unmetamorphosed, struc-

turally simple sedimentary succession consisting of stromatolitic

and evaporitic carbonates with interbedded shales (McArthur and

Nathan groups) overlain unconformably by quartz arenites and

interbedded shales of the Roper Group belonging to the RoperSuperbasin (Jackson et al., 1999), which is now more commonly

referred to as the Beetaloo Basin (Silverman et al., 2007)(Fig. 9B).

The sequence has previously been considered to range in age from

1690 29/25 Ma for the Barney Creek Formation in the McArthur

Group near the base (Page, 1981) to 1429 31 Ma for the McMinn

Formation in the Roper Group near the top (Kralik, 1982), but

organic-rich shales from the Velkerri Formation of the Roper

Group, which occur stratigraphically below the McMinn Formation,

have recently yielded ReeOs dates of 1361 21 Ma and

1417 29 Ma (Kendall et al., 2009). Wells drilled through the Roper

Group succession have encountered widespread live oils and

hydrocarbon shows, including gas seeps, oil staining, uorescence

within drill core, together with widespread occurrence of solid

bitumens formed chie

y by biodegradation of precursor

uid

hydrocarbons (e.g. Muir et al., 1980; Jackson et al., 1986; Powell

et al., 1987; George and Jardine, 1994; George et al., 1994;

Dutkiewicz et al., 2007). Unaltered oil was trapped in quartz-

syntaxial overgrowths during early burial to depths of between 1

and 3 km, while a second oil charge which occurred during deeperburial was subsequently biodegraded to form solid bitumen in the

coarser-grained sandstones (Dutkiewicz et al., 2007). The

Urapunga-4 well (Fig. 9A) drilled in the mid-1980s is particularly

noteworthy because it wasfound to seep live oilfrom thin laminae

within the Velkerri Formation (Jackson et al., 1986; Powell et al.,

1987). The presence of these live oils and hydrocarbon shows

indicate that extensive migration of oil occurred in the McArthur

and Beetaloo basins either during or after the Mesoproterozoic.

Organic-rich sediments occur at ve different stratigraphic

levels in the McArthur Basin and were deposited in a variety of

marine and lacustrine environments. The Mesoproterozoic Velkerri

and overlying Kyalla Shales (in part McMinn equivalent) in the

Beetaloo Basin,are also composed of medium-grey to black,organic

rich, laminatedshale, interbeddedwith thin siltstone andvery ne-grained sandstone deposited under anoixic conditions in shallow-

to moderately deep marine environments (Law et al., 2010). The

TOC content of these shales is typically 2e8% (maximum 12%) and

the organic matter is composed of liquid-prone, type I and type II

kerogen (Warren et al., 1998). The Kyalla and Velkerri Shales in the

Beetaloo Basin range in thermal maturity from immature to over-

mature over a present-day depth range from 350 m to 1500 m,

although the maturity is, in part, due to the effects of Meso-

proterozoic basaltic and dioritic intrusions which are common

within the Roper Group. The maximum cumulative thickness of

these liquid-prone shales is 1600 m. The present-day hydrogen

index (HI) for the middle Velkerri is 281 mg HC/g TOC and the

calculated original HI is 425 mg HC/g TOC, indicating that these

shales have favourable characteristics for the generation of both oil

Figure 6. Geological map and cross-sections through the cupola structure of the Palaeoproterozoic Maksovo shungite deposit, Lake Onega region, northwest Russia (afterMelezhik

et al. 2004). Note that the shungite deposit is conned to layer 6. The thickness of shungite rock decreases towards the margins of the cupola structure (Proles AeB and CeD).

Distribution of the Corg content reveals a mushroom-like structure in the centre of the ore body (Prole EeF).




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and gas (Law et al., 2010). This is conrmed by the fact that cores

taken in them are commonly oil-stained. Rock-Eval pyrolysis Tmaxvalues are c. 480 C near the base of the Kyalla Formation in the

Jamison-1 well at a present day depth of 1700 m (Summons et al.,

1994and eni proprietary data), indicate that the Kyalla Formation

in the centre of the Beetaloo Basin is in the gas generation zone and

suggest that the deeper Velkerri Formation source rocks are likely

to be in the dry gas zone. In contrast, the lower Velkerri Formation

in the Urapunga-4 well to the north has a maximum recorded T maxof c. 460 C at 350 m, indicating a level of thermal maturity

consistent with the late oilewet gas generation zone. The Velkerri

Formation is considered to still retain a large proportion of its

generated hydrocarbons because of its low permeability (Warrenet al., 1998) and it is consequently attracting considerable interest

as a potential unconventional oil and gas play.

Biomarkers (mainly hopanes, monomethylalkanes, alkylcyclo-

hexanes and traces of steranes) in both the oil from the Velkerri

Formation and from marginally mature source rocks in the McAr-

thur Basin indicate that the organic matter is mainly of prokaryotic

(cyanobacterial) origin, possibly with a small contribution from

eukaryotes (Dutkiewicz et al., 2003, 2007). Analyses of the C12hydrocarbon fraction of the oil has revealed a mixture dominated

by n-alkanes of low relative molecular mass without odd-over-even

predominance, 1- and 2-methylalkanes, u-cyclohexyl alkanes and

a series of unresolved mixtures, possibly monomethylalkane

isomers (Jackson et al., 1986;Fig. 9C). The most likely source for the

oils is the organic-rich marine Velkerri Formation, with a possible

component from the slightly older Barney Creek Formation of the

underlying McArthur Group (Dutkiewicz et al., 2007). The most

likely timing of migration of the oils is thought to be after the

emplacement of a dolerite sill (which contains oil inclusions) at c.

1280 Ma and before a phase of signicant Mesoproterozoic struc-

tural inversion and uplift between 1300 and 1000 Ma related to the

assembly of Rodinia (Lindsay, 2002), making these the oldest

known live oils in the world. Their preservation is attributed to the

mild and localised tectonic activity and the lack of metamorphism

experienced by the basin since the source rocks were deposited.

Subtle compositional differences in the content of acyclic and

cyclic biomarkers between the marine Velkerri Formation and the

lacustrine Barney Creek Formation suggest that environmentaldifferences are reected in the composition of organic matter, even

as early as the Mesoproterozoic. All the organic-rich sediments in

the McArthur Basin succession have extremely low abundances of

steranes (biomarkers for eukaryotic organisms) compared to

hopanes. This supports the hypothesis that eukaryotes did not

develop, or at least did not become a volumetrically signicant

component of sedimentary organic material, until later in the

Proterozoic, despite the fact that there are reports of apparent

vertical trace fossil burrows in parts of the Mesoproterozoic Vel-

kerri and Kyalla shales (Law et al., 2010). The Roper Group of

northern Australia (1492 3 Ma:Javaux et al., 2001), and the Ruyan

Group, the age of which is not well constrained but older than

1000 Ma and younger than 1500 Ma: (Xiao et al.,1997) both contain

complex and diverse Mesoproterozoic assemblages of acritarchs.

Figure 7. Location and geology of the Palaeoproterozoic (c. 1.8e2.0 Ga) Mugford Group, northern Labrador, Canada. A. Location map of Labrador. B. Stratigraphy of the Mugford

Group (afterWilton et al., 1993).C. Stratigraphic section of the organic material-bearing horizon, Section4, S.E. Mugford Tickle (afterWilton et al., 1993).D. Geological map of the

Mugford Group showing section locations (afterWilton et al., 1993).E. Thin (2.5 cm thick) coal seaminterbedded with shale, Section4, S.E. Mugford Tickle (afterWilton et al.,

1993). The seam is beneath the pencil. A chert layer forms the lower part of the photograph. F. Thin section (plane polarized light 10) of a thin fractured organic matter layer

(black, upper right) in chert. The organic matter is surrounded by brous quartz veins. The chert contains small ooids with pyrite (opaque) rims (after Wilton et al., 1996).




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4.2. The dullest time in Earths history

Buick et al. (1995)famously described the Mesoproterozoic as

the dullest time in Earths history (p153) and, in a parody of

Winston Churchills statement about the Battleof Britain during the

Second World War, remarked that never in the course of Earths

history did so little happen to so much for so long(p169). Sulphur

isotope compositions of sedimentary pyrites in Mesoproterozoic

black shales from the Vindhyan, Chattisgarh and Cuddapah basins

in India have been interpreted as supporting the existence of

a global sulphidic anoxic ocean at this time, with very low

concentrations of marine sulphate, bacterially reduced in closed

systems (Sarkar et al., 2010). Such extreme environmental condi-

tions would probably have retarded the evolution of multicellular

life and delayed the oxygenation of the biosphere. Analysis of the

d13C patterns in the Mesoproterozoic Bangemall Group of north-

western Australia suggests that the global rate of organic carbon

burial (as a proportion of total carbon burial) and, therefore,

the deposition of potential hydrocarbon source rocks, remained

largely unchanged for nearly 600 million years (1600e1000 Ma)

throughout the Mesoproterozoic (Buick et al., 1995; Xiao et al.,

1997; Brasier and Lindsay, 1998). The static d13C pattern during

the Mesoproterozoic is widely attributed to low bioproductivity

due to a general lack of key nutrients and/or of metabolically

important trace elements such as phosphorus in the Mesoproter-

ozoic oceans (Brasier and Lindsay, 1998;Anbar and Knoll, 2002) as

a result of prolonged tectonic and environmental stability (Buick

et al., 1995).

Figure 8. Optical photomicrographs of selected acritarch taxa from Palaeo- and Mesoproterozoic sediments (from published literature). Scale bars 20mm.a. Ovoidal acritarch from

the Palaeoproterozoic Chuanlingguo Formation, northern China, and diagram showing interpretation of acritarch structure (fromPeng et al., 2009, Fig. 5I,J).b. Spheroidal acritarchfrom the Palaeoproterozoic Chuanlingguo Formation, northern China (fromPeng et al., 2009, Fig. 6A). c.Spiromorpha segmentata(Prasad and Asher) emend (Yin et al., 2005), late

Mesoproterozoic Beidajian Formation, Ruyang Group, Shanxi Province, China (from Yin et al., 2005Fig. 5.1). d. Dictyosphaera delicata (Xing and Liu) Hu and Fu, 1982.Beidajian

formation, Ruyang Group, late Mesoproterozoic, Shanxi Province, China; (i), complete specimen; (ii), enlargement of specimen showing surface reticulate pattern formed by theinterlocking of polygonal plates (from Yin et al., 2005, Fig. 2.4). e . h Tappania plana Yin, 1997. Mesoproterozoic Kamo Group of Central Angara Basin, Siberian Craton (from

Nagovitsin, 2009,Figs. 2c, e).f.Valeria lophostriata(Jankauskas)Jankauskas, 1982. Mesoproterozoic Kamo Group of Central Angara Basin, Siberian Craton (fromNagovitsin, 2009,

Fig. 4e).g.Shuiyousphaeridium macroreticulatum(Yan and Zhou,1992) emend. Yin, 1997, late Mesoproterozoic Beidajian Formation, Ruyang Group, Shanxi Province, China (fromYin

et al., 2005, Fig. 3.6).




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Figure 9. Geology of the McArthur Basin, Northern Australia. A. Simplied geology and locations of some stratigraphic and petroleum exploration wells in the southern McArthur

Basin.B. Generalised stratigraphy showing potential source horizons (S), migrated hydrocarbons (M) and geological ages.C. GCMS Chromatogram of oil recovered from thin laminae

of siltstone enclosed in black mudstone of the Mesoproterozoic Velkerri Formation at a depth of 345.4 to 245.6 m in the Bureau of Mineral Resources (BRM) Urapunga No. 4stratigraphic borehole (afterJackson et al., 1986).




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Despite the apparent low level of bioproductivity at this time,

organic-rich Mesoproterozoic shales frequently contain abundant

sphaeromorphic acritarchs and lamentous sheaths. It has been

suggested that the degree of fossil preservation in some Meso-

proterozoic successions is inversely correlated with the organic

carbon content of the host sediments, as is also the case for the

preservation of acritarchs in many much younger organic rich

sediments (e.g. Vecoli et al., 2009). The shales within the Roper

Group in the McArthur Basin contain abundant and species-rich

assemblages of acritarchs, including distinctive forms belonging

to the species Tappania plana, which have asymmetrically distrib-

uted, hollow, cylindrical processes and bulbous protrusions (Javaux

et al., 2001; Fig. 10). The processes on Tappania have irregular

branching patterns which suggests that it had a cryoskeleton

(Cavalier-Smith, 2002), a characteristic which is considered to be

a fundamental feature of eukaryotic cells (Lamb et al., 2009). There

is a strong environmental control on most Mesoproterozoic

microfossil assemblages with large sphaeromorphic acritarchs (up

to 600 mm) being most abundant in offshore mudstones, while

laments and clusters of small sphaeromorphs characterise tidal

at environments and mixed assemblages occur in shore face and

other shallow water environments. In the Roper Group, the most

abundant and species-rich assemblages occur in estuarine ordeltaic to tide-dominated shoreline facies and there is a clear

onshore-to-offshore pattern of decreasing abundance, declining

and changing dominance of particular species (Javaux et al., 2001;

Fig. 10). The Tappania populations are restricted to distal shelf

shales where they occur with lamentous, probably cyanobacterial,

sheaths and scattered spheroidal acritarchs. The strong facies

control on the distribution of taxa in the Roper Group shows clearly

that environmentally-driven diversication was already well

established in eukaryotic micro-organisms by the early Meso-

proterozoic. However, although the abundance of fossils increases

markedly during the Mesoproterozoic, the morphological diversity

remains relatively low and the similarity and simplicity of the

recorded assemblages still precludes the use of these fossils to

construct age-diagnostic sub-divisions.Well-preserved Proterozoic hydrocarbons provide valuable

information about early biological and biospheric evolution. There

is strong evidence of a global biospheric oxygenation event at c.

1300e1250 Ma in conjunction with a rst-order positive shift in the

marine carbon isotopic record (Altermann, 2004; Schopf, 2004).

This is supported by the appearance of the oldest bedded marine

gypsum deposits and of the earliest, unambiguously multicellular

eukaryotes. This oxygenation event probably played a signicant

role in supporting the more diverse eukaryotic communities

preserved in the Neoproterozoic molecular record and providing

the volume of organic material required to generate commercial

volumes of hydrocarbons.

4.3. The rise and fall of stromatolites..and their role in the worlds

oldest oil and gas elds

Stromatolites (originally called stromatoliths from the Greek

stroma, meaning later and lithos, meaning stone) are dened as

organosedimentary structures produced by the sediment-

trapping, binding and/or precipitation activity of microbial

communities that are dominated by photosynthetic bacteria(McNamara and Awramik, 1994). They are typically nely layered

and mound-shaped. Stromatolites rst appeared at about 3500 Ma,

during the Early Archaean (e.g. Walter et al., 1980). They were

almost certainly the dominant form of life on Earth for much of the

Proterozoic and are, therefore, likely to be a major contributor to

the organic carbon content of any potential hydrocarbon source

rocks deposited during this time. Stromatolites have been reported

from many Archaean sedimentary rocks (Hofmann et al., 1991)

including the 3.43 Ga Strelley Pool Chert in the North Pole Dome

area of the Pilbara Craton of Western Australia (Allwood et al.,

2006), the Tumbiana and Towers formations, also in Western

Australia and the Swaziland Supergroup in South Africa (Byerly

et al., 1986), although the true biological origin of at least some of

these occurrences has sometimes been debated vigorously (e.g.

Buick et al., 1981; Buick, 1991; Lindsay et al. 2005). These early

stromatolites are typically well laminated, simple in form and range

in shape from undulating layered structures to simple domes

(McNamara and Awramik, 1994). The stromatolites in the Strelley

Pool Chert, for example, include simple domes, together with more

complex forms including small, cone-shaped stromatolites

arranged in clusters like egg cartons, some wavy or columnar in

shape and a few branching forms. The existence of several different

forms implies a long evolutionary history and suggests that stro-

matolites may have rst evolved more than four billion years ago

(McNamara, 2009). The environment in which these early Strelley

Pool stromatolites grew is still unclear. The Warraoona Group,

which includes the Strelley Pool Chert, consists mainly of basalts,

several kilometres thick, together with a variety of sedimentary

rocks. Some of the Strelley Pool stromatolites, particularly the egg-

cartonforms, are similar to stromatolites that grow in modern hotspring environments, but it has also been suggested that these very

old stromatolites grew in an extensive shallow sea into which lavas

erupted periodically from nearby volcanoes (McNamara, 2009).

The number of stromatolite taxa increased signicantly during

the Palaeoproterozoic (between 2.5e2.2 Ga and 2.2e1.6 Ga) and

again during the Mesoproterozoic (between 1.65e1.35 Ga and

1.35e1.05 Ga). They were probably at their most diverse during the

late Mesoproterozoic and early Neoproterozoic (Walter and Heys,

1985; Fig. 11) with more than 340 different types having been

identied (McNamara and Awramik, 1994). The comparative rarity

of stromatolites in the Archaean rock record may be more due to

geological, rather than biological, factors (McNamara and Awramik,

1994) since most deposition during the Archaean occurred in

basins within greenstone belts that were characterised by highsedimentation rates. More stable, shallow-water environments that

would have been more conducive to the growth of stromatolites

were probably rare at this time. The increase in the diversity of

stromatolites, and of the microbes that constructed them, in the

rst half of the Proterozoic has been attributed to a decrease in

sediment production rates (Grotzinger, 1990) and/or to the more

widespread development of extensive, stable, shallow water envi-

ronments (McNamara and Awramik, 1994). High sedimentation

rates during the early Proterozoic would have impeded the diver-

sication of stromatolite microstructure by overwhelming biolog-

ical and/or environmental effects as a result of direct precipitation

of laminae or inundation by sediment (Awramik, 1991). Stromato-

lite density (number of stromatolites per unit of rock) was high

during the early Palaeoproterozoic, but diversity (no. of taxa perinterval of time) and abundance (no. of taxa in each basin, summed

for all basins per interval of time) were low (Fig. 11). Individual

stromatolites may have been very abundant at this time, but the

number of different taxa remained low due to the presumed

homogenizing effect of high sedimentation rates.

Conversely, the high diversity of stromatolites during the late

Mesoproterozoic and early Neoproterozoic is interpreted to indi-

cate that sedimentation rates at this time were ideal to promote the

maximum diversity of stromatolite microstructure; sufcient to

allow adequate growth, but not too great to suppress the biological

and/or environmental stimulae for evolution (Grotzinger, 1990).

The distribution of many stromatolite taxa is strongly controlled

by depositional environment (Grey and Thorne, 1985;Fig. 12). For

example, the 2.0 Ga Duck Creek Dolomite in south-western Pilbara,




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contains two different forms of stromatolites. The columnar variety

known as Pilbara is considered to have grown in shallow lagoons

and into the intertidal zone, while the broader-domed branching

variety calledAsperiais interpreted to have grown in pools of water

in high intertidal or supratidal locations. These two distinctive

types of stromatolites (generally referred to as Pilbaraform and

Asperiaform) are recognised in many Proterozoic rocks (McNamara,

2009). Similarly, the proportions and distributions of stromatolitic

and thrombolytic microbialites in the much younger, Neo-

proterozoic (750 Ma) Beck Spring Dolomite in Southern California

are also thought to be controlled primarily by variations in depo-

sitional environment (Harwood and Sumner, 2011).

Figure 10. Palaeobiology of the Mesoproterozoc Roper Group, northern Australia. A. Generalised stratigraphy of the Roper Group showing the radiometric age constraints. B .

Protistan microfossils from the Roper Group. a ,c. Tappania plana, showing asymmetrically distributed processes and bulbous protrusions (arrow in a), b . detail ofa , showingdichotomously branching process.d.Valeria lophostriata.e.Dictyosphaera sp.f.Satka favosa. The scale bar in ais 35mm foraandc; 10mm forb; 100mm ford; 15mm fore; and 40mm

forf.C. Relationship between physical environment, fossil abundance and taxonomic diversity for acritarchs with the Mesoproterozoic Roper Group (afterJavaux et al., 2001).




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It is possible to detect time-related evolutionary changes in

stromatolite morphology that can be used as a basis for stromato-

lite biostratigraphy (e.g. Kumar, 1984) when the environmental

conditions are known and when sufciently rened systematic

descriptions are available. Subdivision and correlation of Precam-brian successions on the basis of their associated stromatolite taxa

has been used widely in the former Soviet Union and rather less so

in China for many years (e.g.Semikhatov, 1976;Zhu, 1982), but it

has been adopted with less enthusiasm elsewhere.

Stromatolite diversity began to decline in stages from the late

Neoproterozoic onwards with marked declines at around 1000 Ma,

at the base of the Cambrian, and again after the early Ordovician.

The initial decline has been attributed to the rise of grazing and

burrowing metazoan (Walter and Heys, 1985). While this may well

have been a factor, as pointed out by Grotzinger (1990), there is an

uncomfortable time gap of, perhaps, 400 million years between the

onset of the decline in stromatolite abundance and the occurrence

of the rst recorded undisputed fossil metazoans. The rate of

stromatolite growth, either through direct precipitation and/or thetrapping and binding of sediment in the microbial mats, ultimately

depends on the rate of supply of carbonate sediment. Proterozoic

seawater may have been highly oversaturated with respect to

calcium carbonate. The appearance of cyanobacterial calcication

during the mid-Neoproterozoic suggests that ocean dissolved

inorganic carbon (DIC) levels had fallen below a critical threshold

by this time, while changes in carbonate cements (e.g. in stro-

matolite microtexture and molar-tooth-structure ) at about the

same time point to a subsequent decrease in the carbonate mineral

saturation state of the ocean (Shields et al., 2009). The decline in

stromotolites from the late Neoproterozoic onwards may also

therefore be, at least in part, related to a decrease in the carbonate

saturation of seawater during the middle and late Proterozoic

leading to a decrease in sediment production and delivery

(Grotzinger, 1990) or, alternatively, to an increase in nutrient levels

in the ocean which is known to have an adverse affect on the

growth of modern stromatolites (McNamara, 2009) or, of course, to

a combination of these factors.

4.3.1. Riphean carbonates, East Siberia (c. 1150e800 Ma)

The Riphean succession on the Siberian Platform and the adja-

cent Yenisey Ridge consists of stromatolitic carbonates and inter-

bedded subordinate clastic sediments ranging in age from c.

1150 Ma to 800 50 Ma (Fig. 13). This succession was deposited

during a period of relative tectonic stability between rifting and

separation of continental blocks from the Siberian Platform in the

Early Riphean (pre-1000 Ma) and collision and later granite intru-

sion during the Late Riphean Baykalian Orogeny (Khain, 1994;

Metelkin et al., 2007; Pisarevsky et al., 2008;Sengor and Natalin,

1996;Vernikovsky et al., 2009.). Broadly age-equivalent deformed

fractured and vuggy Riphean dolostones of the Kamo Group are an

important reservoir in the giant oil and gas elds of the Baykit

Anteclise and on the Siberian Platform (Voronova and Tull, 1993;Fedorov, 1994; Tull, 1997; Postnikov and Postnikova, 2006;

Melnikov et al., 2008;Howard et al., 2009,2012;Fig. 13).

4.3.1.1. Palaeobiology and depositional environment. The Late Mes-

oproterozoic to Early Neoproterozoic succession in eastern Siberia

records the cyclical growth of an extensive carbonate platform,

interrupted by repeated transgressive drowning events during

which the carbonate platform back-stepped and was then smoth-

ered by mainly ne-grained clastic sediment, including turbidites

and organic-rich black shales. Three or four such drowning events

have been recognised within the Upper Riphean succession on the

Yenisey Ridge (Surkov et al., 1996;Fig. 14). Much of the carbonate

platform succession in these 3rd-order cycles consists of biostromal

and biohermal build-ups, up to 60e

70 m thick, consisting of

Figure 11. Variations in density, diversity and abundance of stromatolites and the diversity of acritarchs between the Palaeoproterozoic and the early Phanerozoic (StromatolitesafterWalter and Heys, 1985andGrotzinger, 1990; Acritarchs afterKnoll, 1994andHuntley et al., 2006).




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stromatolites with a wide variety of morphologies including

columnar, domal and laminar forms and dominated by Conophyton,

Colonella, Yakutophyton and Baicalia (CASP, 1998). These are

considered to have been deposited in shallow subtidal to intertidalenvironments, based on comparison with similar Upper Riphean

stromatolite associations in the Finnmark region of Arctic Norway

(Tucker, 1977), although Riphean stromatolites are known to have

occupied a wide range of habitats, including deep subtidal down-

slope environments (Grotzinger, 1989). Colonella is generally

regarded as a subtidal form and other domal, columnar and conical

forms are also probably characteristic of subtidaleintertidal envi-

ronments where stromatolite morphology is controlled principally

by the amount of wave and current energy. The biostromes may

have formed during periods of relatively high energy or when

upward growth of the stromatolites was limited by the availability

of accommodation space. Several of the biohermebiostromal

complexes that crop out on the Yenisey Ridge are capped by

high-energy grainstones and by fractured crystalline dolostoneswhich may represent exposure surfaces, or are directly overlain by

transgressive mudrocks and carbonates belonging to the base of the

next transgressiveeregressive cycle (CASP, 1998).

4.3.1.2. Hydrocarbon source potential and geochemistry.

Riphean organic-rich black shales within the Tungusik Series

(1000e1150 Ma) in East Siberia are widely considered to be one of

the major sources of the oil and gas trapped in fractured and kar-

stied Riphean stromatolitic carbonate reservoirs in the giant elds

of the Baykit Anteclise and the Nepa-Boutuoboya Anticlise on the

West Siberian Platform (Howardet al., 2012; Fig. 14) and may bethe

source of as much as 90% of the known hydrocarbon reserves

(Drobot,1988). The Total Organic Carbon (TOC) content of these late

Mesoproterozoic black shales typically ranges from 0.1 to 11.5%

(Kontorovich,1994; Kontorovich et al.,1996), while the TOCcontent

of the associated stromatolitic carbonates is very low (Surkov et al.,

1996). Riphean successions preserved on the deformed margins of

the East Siberian platform contain thick and laterally extensiveintervals of black shales and carbonaceous mudstones with excel-

lent hydrocarbon source potential (Howard et al., 2012). Potential

source rock intervals include the LowereMiddle Riphean Vedre-

shev and Madra formations (mean TOC of 1.2%), MiddleeLate

Riphean carbonaceous mudstones with TOC values > 10% in the

Cis-Patom Trough (Khabarov, 1995) and thick sequences of black

shale containing source rock intervals with TOC values of 5e10%

and occasionally as high as 20% in the Cis-Enisey Trough

(Voronova and Tull, 1993). The Riphean succession in the eastern

part of the Chunya Basin on the West Siberian Platform contains

a 140 m thick sequence of black carbonaceous mudstone (the Late

Riphean Ayan Formation) with an average TOC content of 1.45%

while the Late Riphean Iremeken Formation in the Yurubchene

Tokhomo zone in the central part of the Baykit Anteclise in WestSiberia contains a 10 m thick unit of carbonaceous mudstone with

an average TOC of 8.27% (Melnikov et al., 2008). The dominant

source of organic-matter in these sediments was probably cyano-

bacteria and acritarchs. Kerogens isolated from the Tungusik Series

black shales have the light (more negative) carbon isotope signa-

ture (d13C PDB of 28 to 32 per mil) characteristic of Proterozoic

organic matter (e.g.Andrusevich et al., 1998;Grantham et al., 1988;

Peters et al., 1995). Crude oils are typically depleted in d13C b y 0e1.5

per mil relative to source rock bitumen (Peters et al., 2005).

Regional palaeogeographic reconstructions suggest that these

Upper Riphean black shales were deposited in a silled-basin,

between a stromatolitic carbonate platform to the east and

a volcanic arc to the west, during repeated periods of transgressive

drowning and backstepping of the carbonate shelf.

Figure 12. Depositional model showing the control of depositional environment on stromatolite forms based on an analysis of the Late Neoproterozoic Buah Carbonates of Oman

(afterCozzi and Alsiyabi, 2004) and the Mezoproterozoic (?) Atar Group Carbonates in the Taoudenni Basin of Mauritania..




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4.3.2. Atar Group, Taoudenni Basin, West Africa (c. 1100 Ma)

Late Mesoproterozoic to Early Neoproterozoic deposition in the

Taoudenni Basin in West Africawas dominated by the deposition of

thick stromatolitic carbonate sequences and associated black shales

belonging to the Atar Group (Trompette, 1973). The Atar Group is

typically around 1000 m thick in the Adrar area on the northern

margin of the basin and thins eastward to 100 m in Algeria and

a fe

Documents

Craig Et Al. 2012 Precambrian Source Rocks JMPG