Université Libre de Bruxelles – Département des Sciences de la Terre et de l’Environnement Musée Royal d’Afrique Centrale – Département de Géologie
The Archaean silicon cycle
Insights from silicon isotopes and Ge/Si ratios in banded iron formations, palaeosols and shales
Camille DELVIGNE
Thèse présentée en vue de l’obtention du grade de Docteur en Sciences
Composition du jury :
Pr. Luc André (MRAC, Belgique) - Promoteur Pr. Nadine Mattielli (ULB, Belgique) - Co-promoteur
Pr. Alain Bernard (ULB, Belgique) - Président Dr. Vinciane Debaille (ULB, Belgique) - Secrétaire
Pr. Marc Chaussidon (CRPG, France)
Pr. Axel Hofmann (Université de Johannesburg, Afrique du Sud)
Septembre 2012
Remerciements
Arrivée au terme de cette thèse, je tiens à remercier les nombreuses personnes qui m’ont aidée à reconstituer le puzzle à l’image mystère que représente une thèse.
Je tiens à adresser mes premiers remerciements à Luc André. Tout d’abord pour m’avoir donné l’opportunité de réaliser ce travail mais surtout pour son enthousiasme inébranlable, sa vision scientifique et ses encouragements.
Pilier essentiel à cette thèse, un grand merci à Damien Cardinal, sans qui de nombreuses pièces du puzzle seraient encore manquantes. Merci pour ta disponibilité de tous les instants, aussi bien pour les questions scientifiques que pour les nombreux jokers « coup de fil à un ami » lors des soucis divers et variés imaginés par le célèbre MC. Et un merci supplémentaire pour être sorti de tes océans pour te frotter aux problèmes de bons vieux cailloux !
J’adresse mes remerciements sincères à Marc Chaussidon et Axel Hofmann d’avoir accepté un aller-retour sur Bruxelles pour évaluer ma thèse. Merci également à Nadine Mattielli, Alain Bernard, Vinciane Debaille d’avoir accepté de juger mon travail.
Cadre incontournable d’une thèse, le labo. Je tiens à remercier Laurence, Nourdine et Jacques pour leur aide quotidienne au labo et lors des analyses ainsi que François, Harold, Pierre-Denis, Frédéric, Claire, Ginnie, Sophie et Katrin pour la vie au labo, leur bonne humeur et les échanges dans un bureau, un couloir ou au téléphone. Merci également à Suzanne pour sa gentillesse et son efficacité. René, pour son aide précieuse dans la préparation des échantillons. Et Didier, pour les jolis posters.
Ensuite, je tiens à remercier en particulier Katrin, Sophie, Jean-Thomas, Ginnie, Eléanore, Vinciane, François et Loïc qui ont apporté une dimension supplémentaire aux conférences.
Merci à nouveau à Axel Hofmann ainsi qu’à Phil Thurston sans qui il me manquerait tout simplement mes échantillons.
Merci également à Nadine, Vinciane, Jeroen et Ivan pour leur gestion du labo MC-ICP-MS et leur aide lors des caprices du MC.
Cette recherche n’aurait pas été possible sans le soutien financier du FNRS que je remercie pour la bourse FRIA qui m’a été accordée pendant quatre années.
Je terminerai par ceux qui me sont le plus chers, ma famille et mes amis.
Merci à mes amis, à ma belle-famille et ma famille pour leurs pensées et leur soutien lors de cette épreuve. Merci en particulier « aux filles », Marjo, Gaëlle, Nath et Marie. Il est bon de vous savoir à l’écoute lors des questionnements et des moments de doutes qui nous ont toutes taraudé l’esprit tôt ou tard et encore.
Merci à mes parents pour leur confiance sans limite dès le début et leur soutien jusqu’à la fin. Merci de m’avoir permis d’arriver jusque là, tout d’abord en éveillant ma curiosité à la moindre fleur ou chenille croisée sur un chemin, ensuite en m’ouvrant les portes de l’université et finalement en préparant de bons petits plats prêts à réchauffer pour la dernière ligne droite.
Et enfin, je ne saurai terminer ces propos sans te remercier, Pascal, toi qui partages avec moi au quotidien les hauts comme les bas. Tu as su à chaque instant avoir les mots qu’il fallait pour me remonter le moral dans les moments de doute et de remise en questions.
Et pour tellement plus, tout simplement, merci.
Dilbeek, H-39
CONTENTS
Abstract... 1
Résumé ... 3
General objectives ... 5
Thesis outline ... 7
CHAPTER 1 ... 9
General overview ... 9
1.1 The history of Earth ... 11
1.1.1 The Hadaean (~4.56-3.8Ga) ... 11
1.1.2 The Archaean (~3.8-2.5 Ga) ... 12
1.1.3 The Proterozoic (~2.5-0.5 Ga)... 13
1.2 The Archaean Earth ... 14
1.2.1 Archaean atmosphere and the “Faint Young Sun” ... 14
1.2.2 Archaean continental crust and its weathering ... 15
1.2.3 Archaean hydrothermalism ... 18
1.2.4 Archaean ocean ... 19
1.3 The Archaean silicon cycle ... 20
1.4 Banded Iron Formations ... 23
1.4.1 Depositional settings ...24
1.4.2 Mineralogies ... 27
1.4.3 Deposition mechanisms... 27
1.4.3.1 Iron oxidation processes ... 28
1.4.3.2 Mechanisms of silica precipitation... 29
1.4.3.3 BIF deposition mechanisms ... 30
1.5 Precambrian cherts ... 34
1.6 Tracers of silicon cycle ... 35
1.6.1 Silicon stable isotopes ... 35
1.6.1.1 Silicon isotopic variations on Earth ... 37
1.6.2 Germanium/Silicon ratio ... 41
1.6.2.1 Ge/Si ratio variations on Earth... 41
1.6.3 Silicon isotopes in BIFs and cherts ... 44
1.6.3.1 Banded iron formations ... 44
1.6.3.2 Precambrian cherts ... 44
References ... 46
CHAPTER 2 ... 59
Analytical developments ... 59
2.1 Outline ...61
2.2 Developments in REE+Y and Ge analysis by HR-ICP-MS ... 62
2.2.1 Sample dissolution ... 62
2.2.2 HR-ICP-MS analysis ... 63
2.2.2.1 Detection limits... 63
2.2.2.2 Correction for interferences on Eu and Gd ... 64
2.2.2.3 Accuracy ... 66
2.3 Controlling the mass bias introduced by anionic and organic matrices in silicon isotopic measurements by MC-ICP-MS* ... 67
2.3.1 Abstract... 67
2.3.2 Introduction ... 68
2.3.3 Material and methods ... 69
2.3.3.1 Instrumentation ... 69
2.3.3.2 Material and sample preparation ... 69
2.3.3.3 Matrix effect counter measures ... 71
2.3.4 Results and discussion ... 73
2.3.4.1 Doped samples ... 75
2.3.4.2 UV treated samples ... 75
2.3.5 Conclusion... 77
References ... 79
CHAPTER 3 ... 80
Desilication in Archaean weathering processes traced by silicon isotopes and Ge/Si ratios ... 80
3.1 Abstract ... 81
3.2 Introduction ... 82
3.3 Methods ... 82
3.4 Palaeosols... 83
3.5 Shales ... 93
3.6 Conclusions ... 97
References ... 98
CHAPTER 4 ... 101
Stratigraphic changes of Ge/Si, REE+Y and silicon isotopes as insights into the deposition of a Mesoarchaean banded iron formation ... 101
4.1 Abstract ... 103
4.2 Introduction ... 103
4.3 Geological setting and sample locations ... 106
4.4 Analytical techniques ... 107
4.5 Results ... 108
4.6 Discussion ... 113
4.6.1 Clastic contamination ... 113
4.6.2 A common parental fluid for Fe- and Si-rich layers ... 114
4.6.3 Constraints on the common parental fluid ... 116
4.6.4 A common siliceous ferric oxyhydroxide precursor ... 118
4.6.5 Sedimentary-diagenetic controls of the temporal δ30Si trend... 119
4.7 Conclusion ... 122
References ... 123
CHAPTER 5 ... 129
Secular changes in the silicon cycle along the Archaean traced by silicon isotopes and Ge/Si ratios ... 129
5.1 Abstract ... 131
5.2 Introduction ... 132
5.3 Materials and methods ... 133
5.3.1 Samples ... 133
5.3.2 Samples preparation and analytical techniques ... 133
5.4 Results ... 136
5.5 Discussion ... 143
5.5.1 Heavier δ30Si values in chert relative to BIF ... 143
5.5.2 Parallel increasing δ30Si trends recorded by BIF and S-chert ... 145
5.5.3 A decreasing trend in Ge/Si ratios recorded by BIF ... 148
5.5.3.1 Changes in ocean inputs ... 149
5.5.3.2 Shifts in δ30Si and Ge/Si values of continent-derived freshwaters... 150
5.6 Conclusions ... 151
References ... 152
CHAPTER 6 ... 157
General conclusions, implications and perspectives... 157
6.1 Insights into continent-derived Si inputs ... 159
6.1.1 Paleoclimatic implications deduced from palaeosols ... 160
6.2 Insights into Si outputs ... 161
6.2.1 Banded iron formations... 161
6.2.1.1 A general model for BIF deposition? ... 163
6.2.1.2 Implications and perspectives for the use of Ge/Si ratio as a source tracer ... 163
6.3 Secular changes in the Si cycle along the Archaean ... 164
6.3.1 Constraining relative role of seawater and high-T fluids as agent of silicification with Ge/Si ratio ... 165
6.4 Crucial need of constrained fractionation factors ... 165
References ... 166
APPENDIX 169
Abstract
The external silicon cycle during the Precambrian (4.5-0.5 Ga) is not well understood despite its key significance to apprehend ancient dynamics at the surface of the Earth. In the absence of silicifying organisms, external silicon cycle dramatically differs from nowadays. Our current understanding of Precambrian oceans is limited to the assumption that silicon concentrations were close to saturation of amorphous silica. This thesis aims to bring new insights to different processes that controlled the geochemical silicon cycle during the Archaean (3.8-2.5 Ga). Bulk rock Ge/Si ratio and Si isotopes (δ30Si) offer ideal tracers to unravel different processes that control the Si cycle given their sensitivity to fractionation under near-surface conditions.
First, this study focuses on Si inputs and outputs to ocean over a limited time period (~2.95 Ga Pongola Supergroup, South Africa) through the study of a palaeosol sequence and a contemporaneous banded iron formation. The palaeosol study offers precious clues in the comprehension of Archaean weathering processes and Si transfer from continent to ocean. Desilication and iron leaching were shown to be the major Archaean weathering processes. The occurrence of weathering residues issued of these processes as major component in fine-grained detrital sedimentary mass (shales) attests that identified weathering processes are widely developed and suggest an important dissolved Si flux from continent to the ocean. In parallel, banded iron formations (BIFs), typically characterised by alternation of iron-rich and silica-rich layers, represent an extraordinary record of the ocean-derived silica precipitation throughout the Precambrian. A detailed study of a 2.95 Ga BIF with excellent stratigraphic constraints identifies a seawater reservoir mixed with significant freshwater and very limited amount of high temperature hydrothermal fluids as the parental water mass from which BIFs precipitated. In addition, the export of silicon promoted by the silicon adsorption onto Fe- oxyhydroxides is evidenced. Then, both Si- and Fe-rich layers of BIFs have a common source water mass and a common siliceous ferric oxyhydroxides precursor. Thus, both palaeosols and BIFs highlight the significance of continental inputs to ocean, generally under- estimated or neglected, as well as the close link between Fe and Si cycles.
In a second time, this study explores secular changes in the Si cycle along the Precambrian. During this timespan, the world ocean underwent a progressive decrease in
hydrothermal inputs and a long-term cooling. Effects of declining temperature over the oceanic Si cycle are highlighted by increasing δ30Si signatures of both chemically precipitated chert and BIF through time within the 3.8-2.5 Ga time interval. Interestingly, Si isotope compositions of BIF are shown to be kept systematically lighter of about 1.5‰
than contemporaneous cherts suggesting that both depositions occurred through different mechanisms. Along with the progressive increase of δ30Si signature, a decrease in Ge/Si ratios is attributed to a decrease in hydrothermal inputs along with the development of large and widespread desilication during continental weathering.
Résumé
Le cycle externe du silicium au précambrien (4.5-0.5 Ga) reste mal compris malgré sa position clé dans la compréhension des processus opérant à la surface de la Terre primitive. En l’absence d’organismes sécrétant un squelette externe en silice, le cycle précambrien du silicium était vraisemblablement très différent de celui que nous connaissons à l’heure actuelle. Notre conception de l’océan archéen est limitée à l’hypothèse d’une concentration en silicium proche de la saturation en silice amorphe.
Cette thèse vise à une meilleure compréhension des processus qui contrôlaient le cycle géochimique externe du silicium à l’archéen (3.8-2.5 Ga). Dans cette optique, le rapport germanium/silicium (Ge/Si) et les isotopes stables du silicium (δ30Si) représentent des traceurs idéaux pour démêler les différents processus contrôlant le cycle du Si.
Dans un premier temps, cette étude se focalise sur les apports et les exports de silicium à l’océan sur une période de temps restreinte (~2.95 Ga Pongola Supergroup, Afrique du Sud) via l’étude d’un paléosol et d’un dépôt sédimentaire de précipitation chimique quasi- contemporain. L’étude du paléosol apporte de précieux indices quant aux processus d’altération archéens et aux transferts de silicium des continents vers l’océan. Ainsi, la désilicification et le lessivage du fer apparaissent comme des processus majeurs de l’altération archéenne. La présence de résidus issus de ces processus d’altération en tant que composants majeurs de dépôts détritiques (shales) atteste de la globalité de ces processus et suggère des flux significatifs en silicium dissout des continents vers l’océan.
En parallèle, les « banded iron formations » (BIFs), caractérisés par une alternance de niveaux riches en fer et en silice, représentent un enregistrement extraordinaire et caractéristique du précambrien de précipitation de silice à partir de l’océan. Une étude détaillée d’un dépôt de BIFs permet d’identifier une contribution importante des eaux douces dans la masse d’eau à partir de laquelle ces roches sont précipitées. Par ailleurs, un mécanisme d’export de silicium via absorption sur des oxyhydroxydes de fer est mis en évidence. Ainsi, les niveaux riches en fer et riche en silice constituant les BIFs auraient une même origine, un réservoir d’eau de mer mélangée avec des eaux douces et une contribution minime de fluides hydrothermaux de haute température, et un même précurseur commun. Dès lors, tant les paléosols que les BIFs mettent en évidence
l’importance des apports continentaux à l’océan, souvent négligés ou sous estimés, ainsi que le lien étroit entre les cycles du fer et du silicium.
Dans un second temps, cette étude explore l’évolution du cycle du silicium au cours du précambrien. Durant cette période, l’océan voit les apports hydrothermaux ainsi que sa température diminuer. Dans l’intervalle de temps 3.8-2.5 Ga, les effets de tels changements sur le cycle du silicium sont marqués par un alourdissement progressif des signatures isotopiques des cherts et des BIFs. Le fort parallélisme entre l’évolution temporelle des compositions isotopiques des deux précipités met en évidence leur origine commune, l’océan. Cependant, les compositions isotopiques des BIFs sont systématiquement plus légères d’environ 1.5‰ que les signatures enregistrées pas les cherts. Cette différence est interprétée comme le reflet de mécanismes de dépôts différents. L’alourdissement progressif des compositions isotopiques concomitant à une diminution des rapports Ge/Si reflètent une diminution des apports hydrothermaux ainsi que la mise en place d’une désilicification de plus en plus importante et/ou généralisée lors de l’altération des continents.
General objectives and thesis outline
General objectives
This thesis aims to bring new insights to different processes that controlled Archaean Si cycle. It addresses the following key issues: (1) Did the Archaean weathering of continents induce a significant desilication? How significant was the dissolved Si riverine outflow to the global ocean? (2) What is (are) the Si source(s) of BIF? What is their deposition mechanism? (3) Did the external Si cycle change along the Archaean? To answer these questions, we have studied (1) a palaeosol (~2.95 Ga) to get insight in Archaean weathering processes and better characterise the potential Archaean Si continental runoff, (2) contemporaneous shales to quantify the relative contributions of primary and weathering-derived minerals into the detrital inputs to the global ocean, (3) contemporaneous BIF with good stratigraphic constraints to assess the Si source(s) and the deposition mechanism of BIF, (4) BIF spanning ages from 3.25 Ga to 2.5 Ga to follow the evolution of the Si cycle using a multi-tracers approach combining silicon isotopes, Ge/Si ratios and rare earth elements (REE).
Thes is outl ine
This thesis is structured into six chapters.
Chapter 1 presents the current knowledge and debates on the early Earth, gives a general overview of Archaean external silicon cycle, reviews the state of the art concerning the two major Archaean Si sinks, banded iron formations and cherts and introduces the silicon isotopes and Ge/Si ratios as tracers of the silicon cycle.
Chapter 2 describes the analytical procedures optimized to provide precise δ30Si compositions, REE+Y and Ge concentrations in different types of samples (mainly banded iron formation, palaeosols and shales).
Chapter 3 focuses on the Archaean weathering processes and the transfer of silicon from continents to ocean. Desilication and iron leaching were shown to be major Archaean weathering processes. The occurrence of weathering residues issued of these processes as major component in fine-grained detrital sedimentary mass (shales) attests that these weathering processes are widely developed and suggests an important dissolved Si flux from the growing continents to the ocean.
Chapter 4 proposed a deposition mechanism for a 2.95 Ga banded iron formation. A common precursor for silica-rich and iron-rich layers through the silicon adsorption onto Fe-oxyhydroxides forming a siliceous ferric oxyhydroxides precursor is evidenced.
Besides, the parental water mass from which banded iron formation precipitated is identified as a seawater reservoir mixed with significant freshwater component and very limited amount of high temperature hydrothermal fluids.
Chapter 5 explores secular changes in the Si cycle along the Precambrian. The progressive increase in δ30Si values along with decreasing Ge/Si ratios recorded by BIFs and cherts within the 3.8-2.5 Ga time interval highlights the decrease in seawater temperature and the relative increase in continental inputs. Besides, the systematic lighter δ30Si signature recorded by BIFs compared to cherts is discussed in term of different deposition mechanisms.
Chapter 6 summarises the main results and conclusions of this work and discusses the perspectives open by this study.
CHAPTER 1
G ENERAL OVERVIEW
CHAPTER 1
General overview
1.1 The history of Earth
Typically the history of Earth is divided into two supereons: the Phanaerozoic, spanning ages from today back to 500 Ma ago, and the Precambrian that covers time from 500 Ma ago back to the Earth’s formation, about 4.56 Ga. The Precambrian, which lasts about 4.0 Ga, is subdivided into three eons: the Hadaean (~4.5-4.0 Ga), the Archaean (~4.0-2.5 Ga) and the Proterozoic (~2.5-0.5 Ga). Let us describe them in a nutshell.
1.1.1 The Hadaean (~4.56-3.8Ga)
The Hadaean begins with the Earth accretion about 4.56 Ga (Patterson, 1956) and ends with the first oldest dated rocks (~3.8 Ga). In the absence of rock records on Earth, most of our knowledge of this time span comes from studies on meteorites, the Luna as well as physical and geochemical models.
The Earth was formed from the solar nebula about 4.56 Ga ago. Only about 11 Ma after the formation of the solar system, the Earth achieved 99% of its current mass and already had segregated its core (e.g., Kleine et al., 2002; Yin et al., 2002). The formation of the moon by giant impact occurred appoximately 35 Ma later (Halliday, 2004). It is of common view to imagine the Hadaean Earth as covered by a magma ocean. From rare sedimentary zircons relicts (4.4-4.0 Ga), authors suggested the presence of a significant quantity of water and a transitory proto-crust (e.g., Mojzsis et al., 2001; Caro et al., 2005;
Bibikova, 2010). This suggests that temperatures were quickly low enough to allow the existence of liquid water and that the magmatic ocean cooled. Then, the Earth with a core, a mantle, oceans and an atmosphere would have already existed only 150 Ma after the Earth accretion. However, Hadaean proto-crust, oceans, atmosphere were not comparable to today’s. The end of the Hadaean is marked by an intense bombardment from meteorites called “Late Heavy Bombardment” (Gomes et al., 2005) destroying any rocks and sterilizing the Earth of any life that could have already emerged.
1.1.2 The Archaean (~3.8-2.5 Ga)
The Archaean period streching from 3.8 to 2.5 Ga was a key time in building-up the modern Earth. This crucial period in our Earth’s history is marked by fundamentals changes: (1) the onset of a plate tectonic; (2) the growth of continents and (3) the emergence of life.
As a consequence of an intense heat production, the Archaean plate tectonic was more complex than the current one, with numerous small plates moving quite fastly. To this horizontal tectonic, a vertical tectonic driven by gravity probably surperposed (e.g., Hickman, 2004; Van Kranendonk, 2007). Although how and when tectonic occured remain hotly controversial, there is a growing concessus that a form of tectonic operated on a global scale by 2.8 Ga ago (e.g., Van Kranendonk, 2004; Smithies et al., 2005). With the end of intense bombardment at 3.8 Ga, the newly formed crust survived, stabilised and continents appeared. It has been suggested that a first supercontinent, called Vaalbara, existed around 3.3 Ga ago (e.g., de Kock et al., 2009 for a recent example).
Earth was mostly covered with anoxic and acidic oceans where first life appeared and evolved. The first photosynthetic organism appeared probably 3.5 Ga ago as evidenced by controversed stromatolites-like structures (see Precambrian Research, v. 158, 2007).
Dating the oldest life form is intricate but it is generally accepted that first life forms are older than photosynthetic life, perhaps many hundreds of millions of years older (Schopf, 1992). Colonies of cyanobacteria (forming stromatolites) producing oxygen through photosynthesis are regarded as the builders of our oxygenated atmosphere. However, despite cyanobacteria started to produce oxygen about 3.5 Ga ago, atmosphere remained anoxic for a long time as Earth crust, reduced volcanic gasses and oceans acted as oxygen sinks. Main evidences pointing to low oxygen levels in Archaean atmosphere are (1) the occurrence of redox-sensitive minerals such as uraninite, pyrite, siderite in Archaean sediments (e.g., Ono et al., 2000; Frimmel, 2005; Hofmann et al., 2009) ; (2) the Fe-mobility in palaeosols formed prior to 2.2 Ga (see review of Rye and Holland, 1998); (3) the preservation of large mass independent fractionation of sulfur isotopes conditioned by the absence of ozone protection from UV (e.g., Farquhar et al., 2000, 2003; Bekker et al., 2004). As this study focuses on this period, a more detailed state of knowledge will be addressed in section 1.2 of this thesis.
1.1.3 The Proterozoic (~2.5-0.5 Ga)
The Proterozoic, which lasted from about 2.5 Ga to 0.5 Ga ago, was the period of early life development and explosion. Proterozoic is characterised by three main features (1) a modern-like plate tectonic; (2) the rise of oxygen in the atmosphere and (3) the development and evolution of early life.
Despite having a modern-like dynamic, the movement of plates were faster than today’s as a consequence of a hotter magma. Then, frequent collisions led to the building-up of large continents and the formation of a super-continent called “Rodinia” about 1.2-1.1 Ga ago (e.g., Piper, 1976, 2000; Dalziel, 1991; Meert and Powel, 2001). Parts of this super- continent survived and are now pieces of North America, Australia, Western Africa and Antarctica.
When the various oxygen sinks (oceans, minerals, volcanic gasses) were saturated, the photosynthetically produced oxygen begun to accumulate in the atmosphere. The rise of oxygen in atmosphere about 2.4-2.3 Ga is called the “Great Oxydation Event” (GOE) (Fig.
1.1). It marks the transition between the late Archaean and the early Proterozoic. Besides production of oxygen through photosynthesis was the primary driver of oxydation during the GOE, a number of other processes such as changes in the redox potential of volcanic gases are postulated to have played a role in changing the redox state of the ocean- atmosphere system (see reviews of Holland, 2009 and Pufhal and Hiatt, 2012).
Fig. 1.1 Prevailing view of atmospheric oxygen evolution over time. The large increase at 2.4 Ga is commonly known as the “Great Oxydation Event”. From Kump, 2008.
Produced oxygen acted as a toxic element for anaerobic life that was wiped out forcing the evolution through aerobic life. During the Proterozoic, life evolved from anaerobic life into aerobic life, eucaryotes, and multicellulars.
1.2 The Archaean Earth
Understanding Archaean surface environments has been challenging for decades. In order to address this issue, Archaean oceans chemistry offers useful window into past processes as they intimately link atmosphere, continents and hydrothermal fluids. The next sections describe our current state of knowledge on these Archaean reservoirs.
1.2.1 Archaean atmosphere and the “Faint Young Sun”
Based on observations of solar-like stars and modeling, the solar luminosity was likely reduced by 30% 4.6 Ga ago (Gough, 1981; Sagan and Mullen, 1972). If the composition of the atmosphere had remained unchanged, Earth’s surface should have been frozen before ~2.2 Ga (Sagan and Mullen, 1972; Kasting et al., 1988). However, many evidences suggested the existence of liquid water on Earth’s surface at a very early stage in its history (e.g., Valley et al., 2002; Mojzsis et al., 2001; Nutman, 2006), perhaps as early as the late stages of accretion. This apparent contradiction between solar evolution models and geological evidence for liquid water has been called the “Faint Young Sun” paradox.
The most widely accepted theory to solve this problem is the presence of enhanced concentrations of greenhouse gases in early atmosphere keeping the Earth warm. CO2
and methane (CH4) are the more likely candidates for such greenhouse gasses. However, the combination and relative abundances of these greenhouse gasses are under debate (see Fig. 1.2 for CH4 and CO2 levels proposed by Lowe and Tice, 2007). Models based on CO2 alone require CO2 concentrations well above CO2 levels deduced from sediments data (Hessler et al., 2004; Rye et al., 1995; Rosing et al., 2010). Moreover, this was challenged by Sleep and Zahnle (2001) who argued that CO2 was removed from the Archean atmosphere-ocean system by carbonitization of seafloor. Then, a combined effect of CO2
and CH4 was envisaged (Rye et al., 1995; Pavlov et al., 2000; Lowe and Tice, 2004; Kasting, 2005). This was based on the premise, now widely accepted, that levels of atmospheric oxygen were low before 2.4 Ga ago. When atmospheric O2 rose, the concentrations of
CH4 and other reduced gases should have decreased dramatically, possibly triggering the observed glaciations (e.g., Evans et al., 1997). However, it should be emphasised that the
“Faint Young Sun” problem is not yet solve and other parameters such as a lower albedo contributed to keep Earth’s climate clement. Taking into account of such parameters in models alleviate the need for extreme greenhouse gasses concentrations to satisfy the Faint Young Sun paradox (von Paris et al., 2008; Rosing et al., 2010). Lowe and Tice (2004, 2007 and references therein) evidence a close link between crustal, atmospheric, climatic and biological evolutionand propose a tectonic control.
Fig. 1.2 Schematic model of fluctuations of atmospheric pCO2 and pCH4 between 3.5 Ga and 2.1 Ga suggested by Lowe and Tice (2007).
1.2.2 Archaean continental crust a nd its weathering
The Tonalite–Trondhjemite–Granodiorite (TTG) suite composes up to 90% of the juvenile Archaean continental crust (Martin et al., 2005). These TTG are closely associated to greenstone belts made up of komatiites, basalts and sediments. The formation of TTG and komatiites are of strong interest to comprehend the Archaean Earth as their formation is restricted to the Archaean. Processes producing TTG originate in a tectonic dynamic operating with important heat flux where the subducted oceanic crust reached melting temperatures in contrast to modern tectonic where it dehydrates inducing the melting of the above mantle wedge (Martin, 1986). Besides displaying probably different composition than the modern continental crust, the continental landmass was
considerably smaller at this epoch. Many models, highly debated, have been proposed for the growth of continental crust over time (e.g., Reymer and Schubert, 1984; Collerson and Kamber, 1999; Condie, 1998, 2000). As illustrated by figure 1.3, the estimated volume of the Archaean continents greatly varies depending on the model. As these models are controversial, we’ll satisfy here to state that continental inputs to oceans would therefore be considerably smaller than nowadays. However, such low continental fluxes might have been compensated by an intense source-rock weathering (e.g., Fedo et al., 1996; Sugitani et al., 1996; Lowe and Tice, 2004; Hessler and Lowe, 2006).
Fig. 1.3 Different models for the growth of continental crust over time. From Lowe and Tice, 2007.
Ancient weathering processes are not well understood because ancient weathering profiles (palaeosols) are scarce and most of them underwent diagenesis and metamorphism that changed their mineralogy. Many studies report intense source-rock weathering during Archaean (e.g., Fedo et al., 1996; Hessler and Lowe, 2006). In the absence of biological effects, the factors invoked to explain intense weathering include heavy rainfall, elevated surface temperatures and higher atmospheric pCO2 (e.g., Lowe and Tice, 2004; Hessler and Lowe, 2006). Discriminating between temperature and pCO2
as the principal factor driving weathering is not yet possible and most probably both played significant roles (Hessler and Lowe, 2006). Surface conditions on early Earth are poorly constrained and, despite extensive studies, uncertainties remain.
Palaeotemperature models yield large discrepancies of the results depending on the approach implemented. Oxygen isotopic composition of cherts (and recently supported
by silicon isotopes) provide evidence for a hot Archaean ocean (70±15°C) (Knauth and Lowe, 2003; Knauth, 2005; Robert and Chaussidon, 2006). High chemical weathering indices are generally interpreted as a result of warm climatic conditions although a cause- and-effect link cannot be clearly established. However, Hofmann (2005) warned that the interpretation of weathering indices requires caution as post-depositional effects such as hydrothermal alteration may distort weathering indices. Although not precluding a hot climate, a study based on quartz weathering features moderated that Archaean climate was not necessarily hot and may be consistent with modern weathering temperature (Sleep and Hessler, 2006).
Besides elevated surface temperature and high pCO2, the low atmospheric pO2 induced deep changes in the weathering of redox-sensitive minerals. Significant loss of Fe is common in Precambrian palaeosols (Rye and Holland, 1998). This contrasts with constant Fe concentrations in modern weathering profiles formed under oxic conditions, where dissolved Fe2+ is oxidized to Fe3+ and remains as Fe(III)-oxides in the profile (Maynard, 1992). It is generally considered that Fe loss is typical of subaerial weathering at low pO2
prior to 2.2 Ga(Rye and Holland, 1998). However, Fe-depletion should not be considered as a typical feature of palaeosol as diagenetic processes can mimic this Fe depletion even under an oxidizing atmosphere (Ohmoto et al., 1996). In contrast, redox-sensitive minerals in a reduced state (e.g., pyrite, uraninite, siderite) remain stable under Archaean atmosphere with low pO2. This contrasts with modern oxidative weathering considered as one of the most important sulfate sources to modern oceans. This input of sulfate to oceans likely not operated during Archaean time, leading to low sulfate contents in the Archaean oceans.
To sum up, Archaean continental inputs to oceans differ from nowadays through (1) intense weathering producing important flux of labile elements, typically Ca, Na, Sr; (2) reducing conditions that have promoted large Fe soils outflows but low sulfate fluxes.
However, riverine supplies to the global ocean might have been partly counterbalanced by the smaller size of the continental landmass (e.g., De Wit and Hart, 1993; Collerson and Kamber, 1999; Condie, 2000).
1.2.3 Archaean hydrotherma lism
As nowadays the excessive heat of the mantle must have been dissipated through conduction, convection associated to the generation and subduction of oceanic crust and, convection of seawater into the oceanic crust. In response to greater heat fluxes, the total length of medio-oceanic rides (MORs) might have been twice as long as today (Isley, 1995). This might have culminated in an order of magnitude larger hydrothermal activity in the Archaean (De Wit and Hart, 1993). Although more vigorous, sea-floor hydrothermal systems appear to have operated in a relatively constant way throughout the Earth's history: early to mid-Archaean environments can be interpreted in terms of present-day settings and chemistry (e.g., de Ronde et al., 1994, 1997; de Vries and Touret, 2007; Hofmann and Harris, 2008). Then, most models rely on the assumption that chemical properties of Archaean hydrothermal fluids were similar to those of modern hydrothermal fluids. However, recent thermodynamic models pointed that this might not be valid (Shibuya et al., 2010; Wang et al., 2009). Shibuya and coworkers (2010) predicted that if the Archaean seawater was slightly acidic, CO2-rich and SO4-poor, high temperature alteration processes of the oceanic crust would have led to the generation of highly alkaline (pH>10), Si-enriched and Fe-depleted hydrothermal fluids. In contrast Wang et al.
(2009) pointed out that alkaline, Si-Fe-enriched hydrothermal fluids can be merely formed from the hydrothermal leaching of komatiites.
Except Shibuya et al. (2010), most studies suggest that because of reducing conditions and low sulfate concentrations of Archaean seawater, the hydrothermal delivery of Fe to the ocean was higher (Walker and Brimblecombe, 1985; Isley, 1995; Canfield, 1998; Kump and Seyfried, 2005) even under relatively low-temperature conditions (Hofmann and Harris, 2008). Therefore, hydrothermalism must have represented an important Fe flux to Archaean oceans where Fe must have been accumulated in response to the reducing conditions of the deep ocean.
1.2.4 Archaean ocean
From previous sections, it appears that many parameters controlling the chemistry of the ocean were dramatically different during the Archaean time span. Although not exhaustive, following dramatic changes between modern and Archaean conditions can be cited: (1) Archaean atmosphere was CO2-rich and O2-poor (see section 1.2.1); (2) surface temperatures were higher increasing solubility of elements (see section 1.2.1); (3) the continental fluxes were smaller (see section 1.2.2); (4) the mantle-derived inputs were likely more important (see section 1.2.3); (5) chemistry of inputs to oceans evolved through time (see section 1.2.2 and 1.2.3); (6) biogenic processes were not yet operating, or at least in a different way. Therefore, it seems difficult to imagine how Archaean ocean may not have been very different in composition from the modern ocean. In broad outline, following assumptions can be made.
The relative balance between mantle-derived vs. continental-derived inputs has been the subject of much debate. Veizer et al. (1982) popularized the view that Archaean balance between continental and hydrothermal inputs could have been reversed relative to present-day with Archaean ocean dominated by hydrothermal inputs. This sounds consistent with on the one hand, reduced continental inputs reflecting the smaller continental volume at the time and, on the other hand, higher hydrothermal fluxes related to greater mantle heat flow. Numbers of studies attempted to better quantify the relative contributions using geochemical and isotopic proxies (e.g., Derry and Jacobsen, 1990; Bau et al., 1997). Recently, some authors argued that the strong dominance of hydrothermal inputs into Archaean ocean cannot be simply explained by increased Archaean hydrothermal activity and reduced continental volume (Kamber and Webb, 2001). Taking into account reduced continental inputs, they calculated that hydrothermal flux should have been ten times greater than today’s. Considering such an increase as unrealistic, Kamber and Webb (2001) further invoked a reduced erosion rate in addition to the smaller continental volume compared to nowadays.
Enrichment of CO2 in the atmosphere arguably had pronounced effects on the hydrosphere because high pCO2 tends to lower the pH of solutions in contact with the atmosphere. Some studies proposed that CO2 concentrations of Archaean atmosphere might have been up to 300 times the Present Atmosphere Level (PAL) (Kasting, 1993;
Lowe and Tice, 2004). Such high CO2 levels in the atmosphere would have induced a pH of the oceans between 5.4 and 8.6 (Grotzinger and Kasting, 1993). However, debates remain on how high were atmospheric CO2 levels (Kasting, 2005; Lowe and Tice, 2004; Sumner and Grotzinger, 2004; Catling et al., 2001).
As the level of oxygen in atmosphere has risen over time, it is expected that the oceans too will encompass a progressive oxidation. Then, during Archaean oceans are considered to be at a reduced state in response to low atmospheric pO2. However, debate remains on the possible existence of a redox-stratified ocean where surface waters could have been oxidized. A lack of oxygen, at least in watermass below the hypothetic oxycline, allows iron to remain in its reduced state. Combined with very low levels of sulfates (very small inputs from rivers in the absence of oxidative weathering of sulfur), great amounts of iron tend to accumulate in Archaean oceans.
The most widely accepted view is to assume that the Archaean oceans were anoxic and acidic with lower carbonate and sulfate contents but higher concentrations of Ca, Fe, Ba, Si, Na, Cl compared to modern oceans.
1.3 The Archaean silicon cycle
Our current understanding of Precambrian oceans is limited to the assumption that Si concentrations were close to saturation of amorphous silica (Siever, 1992). Due to the strong temperature-dependent solubility of silica in pure water (Fig. 1.4) (Gunnarsson and Arnorsson, 2000) and the absence of a consensus on temperature prevailing in Archaean times, Si concentrations of Archaean oceans remain poorly constrained. Depending on assumed temperature, Si concentrations estimates range between 30 and 60 ppm (e.g., Morris, 1993; Siever, 1992), which implies a long residence time on the order of 105 years in the Precambrian ocean (Siever, 1992).
Fig. 1.4 Temperature-dependent solubility curves of quartz and amorphous silica at Psat. Calculated from Gunnarsson and Arnorsson, 2000.
The origin of Si in the Precambrian ocean is controversial (Maliva et al., 2005). As nowadays Si inputs were likely from hydrothermal fluids and continental-derived freshwaters. Besides a presumably reverse balance between hydrothermal and continental fluxes during the Precambrian (e.g., Bau and Möller, 1993; Bau et al., 1997;
Kamber and Webb, 2001), Si concentrations of both fluxes may have been different as a result of contrasted environments.
The Si concentration of modern hydrothermal fluids along modern mid-oceanic spreading centres is generally considered to be controlled by equilibrium with quartz at temperatures ~200-400°C and pressures of 100-500 bars in hydrothermal convection cells (Von Damm et al., 1985, 1991). As various pressure and temperature conditions are encountered, hydrothermal fluids may thus display different Si concentrations. Modern hydrothermal fluids have Si concentrations ranging from ~450 to ~650 ppm (Mortlock et al., 1993). According to recent studies, silicon in Archaean hydrothermal fluids can theorically reach concentration as high as 1680 ppm or 3000 ppm, which is at least 4 to 7 times higher than many modern hydrothermal fluids (Wang et al., 2009; Shibuya et al., 2010; see section 1.2.3).
River influx of dissolved silicon to the Archaean oceans would strongly depend on the surface of emerged continents exposed to weathering and the weathering regime. Most
of crustal growth models agree with a smaller volume of continental crust during Archaean compared to nowadays. Although weathering regimes might have been more aggressive (Fedo et al., 1996; Sugitani et al., 1996; Lowe and Tice, 2004; Hessler and Lowe, 2006), potentially increasing the dissolved Si load in rivers, it seems reasonable to assume that Si continental inputs were probably by far less than present-day due to small volume of continental crust.
However, despite variances affecting Si inputs, the pivotal change between modern and Archaean Si cycles are the Si outputs. In the absence of biologically-mediated precipitation, the most likely Si removal processes are (1) the direct precipitation of silica minerals and (2) the pervasive silicification of volcanoclastic sediments and their crystalline seafloor basements. Sorption of silicon on clay minerals, Fe oxides, or organic matter could also have played a role in nucleation and/or precipitation (e.g., Perry and Lefticariu, 2003; Siever, 1992; Fischer and Knoll, 2009). Then, cherts formed either by direct precipitation (C-cherts) or silicification of a rock precursor (S-cherts) as well as cherts of banded iron formations (BIF) likely represent dominant sinks of silica in Precambrian oceans. However, for both cherts and BIFs, the source(s) of Si (continental- derived freshwaters, hydrothermal fluids, seawater or a mixture of these reservoirs) as well as their precipitation mechanisms remain unclear. Therefore the impact of their deposition on the silicon cycle is not fully understood.
1.4 Banded Iron Formations
Banded iron formations (BIFs) are chemical precipitates characterised by the alternation of iron-rich and silica-rich layers with three scales of banding: microbands (<1mm), mesobands (~1mm – 10cm) and macrobands (>1m) (Trendall and Blockley, 1970) (Fig. 1.5).
According to the definition of James (1954), BIF typically BIF contain more than 15 wt.% of iron.
Fig. 1.5 Illustrations of the BIF banding scales. A- Outcrop of the ~2.95 Ga Mozaan BIF in the White Umfolozi Inlier (South Africa) where the macrobanding (~1m) is visible. B- Sample from the ~2.95 Ga Witwatersrand Supergroup with alternating cm-scale layers of jasper (red layers) and Fe-oxides layers (black layers); C- Sample from the ~2.5 Ga Transvaal Supergroup with alternating cm- and mm-scale layers
of cherts (white layers) and Fe-oxides layers.
BIFs are largely restricted to the Precambrian with most occurrences ranging between 3.8 Ga and 1.8 Ga (Fig. 1.6). The peak in BIF deposition about ~2.5 Ga appears to correlate with major changes in the Earth’s history such as the rise of atmospheric oxygen and the change from anoxic to oxic conditions in the ocean (Canfield, 2005; Holland, 2006) (Fig.
1.6).
Fig. 1.6 Schematic diagram showing the relative abundance of Precambrian BIFs through time with calculated curves for the atmospheric evolution of oxygen and carbon dioxide (Kasting,2001, 2004;
Kasting and Catling, 2003; Pavlov and Kasting, 2002). Estimated abundances are relative to the Hamersley Group BIF volume taken as a maximum. From Klein, 2005.
Their abundance in the Archaean/early Proterozoic eons and their absence thereafter clearly reveal different Si and Fe cycles before 1.8 Ga ago. As BIFs represent an important sink of Si during this time span, it is of great interest to better understand the BIF deposition mechanism which has not yet reach a consensus (see section 1.4.3). Before exploring different BIF deposition mechanisms that have been proposed, a brief look at deposition settings and mineralogies is instructive.
1.4.1 Depositional settings
Based on the BIF size and lithologic associations, two type of BIF are distinguished:
Algoma- and Superior-type BIF (Gross, 1965). Algoma-type BIFs are relatively small, associated with volcanogenic rocks and hosted in greenstone belts. Typical lateral extensions are less than 10 km and thickness is less than 50 m. However, these characteristics do not indicate that all Algoma-type BIF were originally small, as most have been affected by deformation and tectonic dismemberment, implying that their true size and extent are likely underestimated in global compilations of BIF (Gole and Klein, 1981).
Reconstruction of the original depositional settings very difficult but favored depositional environments for this type of BIF include island arc/back arc basin (Veizer, 1983) and intracratonic rift zone (Gross, 1983). BIF of Isua (~3.8 Ga, West Greenland) and Barberton (~3.5-3.1 Ga, South Africa) greenstones belts are typically representative of Algoma-type BIF. Superior-type BIFs are larger and associated with other sedimentary units. Several
BIFs classified as Superior-type BIF have extension over 105 km² (Beukes, 1973).
Deposition occurred in relatively shallow marine conditions under transgressing seas, perhaps on the continental shelves of passive margins (Gross, 1965). Superior-type sequences include the Hamersley Group (~2.5 Ga, Western Australia), the Pongola, Witwatersrand and Transvaal Supergroups (~2.5 Ga, South Africa). Note that a third type, the Rapitan-type, was defined for Neoproterozoic BIFs associated to glacio-marine deposits and resulting from anoxic conditions caused by an ice-covered ocean, the
“Snowball Earth” (Gross, 1973). As this thesis focus on the Archaean, this type will not be considered further.
Fig. 1.7 Secular distributions of mantle superplume breakout events (Abbott and Isley, 2002) and BIF deposits. (From Bekker et al., 2010)
Isley (1995) and Isley and Abbot (1999) highlighted a correlation between major BIF deposition events and global plume events, suggesting a causal link between both (Fig.
1.7). Mantle superplume volcanism may promote BIF deposition by (1) increasing the Fe flux through continental weathering and/or through submarine hydrothermal processes, (2) increasing the number of tectonic environments appropriate for BIF deposition, (3) promoting global anoxic, Fe-rich hydrothermal plumes in the shallow to intermediate marine water column (Isley and Abbot, 1999). However, Huston and Logan (2004) pointed out that the relation between Superior-type BIFs and global mantle plume events is not direct. Then, Algoma-type BIFs were regarded to reflect intrabasinal pulses of magmatic and hydrothermal activity during the deposition of volcano-sedimentary
greenstone successions. Superior-type BIFs are considered to have formed on continental shelves during periods of global high sea levels and pulses of hydrothermal activity. If this model is correct, then a distinction between Superior- and Algoma-types of BIFs has some merit. Indeed, the geochemistry of Algoma-type BIFs may reflect local hydrothermal conditions whereas Superior-type BIFs may record the large-scale chemistry of the oceans during their formation, although a potential influence by nearby is not excluded.
Although in some cases it is difficult to resolve whether a BIF belongs to Algoma- or Superior-type, Algoma-type BIFs are more common in Archaean (>2.5 Ga) whereas Superior-type BIFs appeared in the record at ca. 3.3 Ga but become significant at about 2.6 Ga ago (Fig. 1.8).
Fig. 1.8 Variations in the abundance of Algoma- and Superior-type BIF through time (from Huston and Logan, 2004).
The secular changes in the style of BIF deposition may reflect a higher mantle heat flux and a limited occurrence of continental shelves during Archaean favoring Algoma-type BIF. Increasing number of continental shelves after a major plume event at 2.75-2.70 Ga that lead to important continental crust formation allowed the widespread deposition of Superior-type BIF (Clout and Simonson, 2005; Bekker et al., 2010). There is a global consensus to attribute the end of BIF deposition at about 1.8 Ga as a consequence of the oxygenation of the atmosphere (e.g., Huston and Logan, 2004).
1.4.2 Mineralogies
Mineralogically, Algoma- and Superior-type BIF are similar. Most BIF have undergone various grades of metamorphism from the late-diagenetic to the amphibolites facies metamorphic grades. Major mineral phases in late-diagenetic and low-grade metamorphic BIFs are magnetite, hematite, pyrite, greenalite, stipnomelane and minnesotaite and carbonates including siderite and members of dolomite-ankerite series (e.g., Klein, 2005). Chert (mainly microcrystalline quartz) is ubiquitous and is almost always present. Depending on the major Fe-bearing mineral phase, iron formations are classified as oxide, carbonate, silicate and sulfide facies with several mixed facies (James, 1954; Klein and Beukes, 1993). The sulfide facies should better be considered as pyritic carbonaceous shale or slate than BIF (Bekker et al., 2010). Original precipitates of BIFs have not yet been unequivocally identified. One the one hand, some consider that chert layers formed by recrystallisation of an amorphous silica precursor (e.g., Hamade et al., 2003; Maliva et al., 2005; Posth et al., 2008; Steinhoefel et al., 2009; Wang et al., 2009).
On the other hand is the hypothesis of a common precursor such as a siliceous ferric oxyhydroxide gel or Al-poor hydrous iron silicate mud (e.g., Lascelles, 2007; Fischer and Knoll, 2009).
1.4.3 Deposition mecha nisms
Banded iron formations represent a style of sedimentation for which clear modern analogues do not exist, at least at a similar scale. The controls on the deposition of BIFs have long been contentious and an accepted global theory of BIF deposition is still lacking. However, it is also unclear whetever a unique theory would accomodate the diversity of BIF as their proposed deposition settings as well as mineralogies are varied.
Controversy still surrounds: (1) the processes of iron precipitation; (2) the origin of chert;
and (3) depositional systems.
1.4.3.1 Iron oxidation processes
Most studies on BIFs have focused on the origin of Fe-rich bands. It is generally accepted that Fe2+ was sourced from hydrothermal alteration of oceanic crust (e.g., Jacobsen and Pimentel-Klose, 1988; Derry and Jacobsen, 1990; Bau and Möller, 1993). Another common assumption is that the oceans (at least at depth) were anoxic and therefore capable of transporting and concentrating Fe2+ dissolved in seawater (e.g., Cloud, 1968; Holland, 1973; Drever, 1974). In addition to low oxidizing potential and high hydrothermal iron flux, low marine sulfate and sulfide concentrations (Habicht et al., 2002) allowed an amplified marine reservoir of dissoved iron. Proposed mechanisms of oxidation of Fe2+ and subsequent precipitation of ferric iron are still debated (e.g., Ohmoto et al., 2006; Beukes and Gutzmer, 2008; Bekker et al., 2010). Three main models have been proposed and are briefly described below.
Oxidation of Fe(II) by cyanobacterially-generated O2
This widely quoted model involves the inorganic oxidation of dissolved Fe(II) with photosynthetically procuded oxygen by cyanobacteria (Cloud, 1965, 1973):
4Fe2+ +O2 + 10H2O 4Fe(OH)3 + 8H+
Under an anoxic atmosphere, the produced oxygen would have confined to localised
“oxygen oases” associated with cyanobacterial blooms or formed a stratified ocean with a thin oxic zone overlying an anoxic water column (e.g., Klein and Beukes, 1989). Main rebuttal against this hypothesis is that most Archaean and early Paleoproterozoic hydrogeneous sediments (e.g., carbonates and BIFs) do not show geochemical evidence of a distinct redoxcline even in shallow environments (e.g., Kamber et al., 2004;
Alexander et al., 2008; Planavsky et al., 2010). This has recently been used to question the iron oxidation through mixing of oxic and anoxic waters (e.g., Planavsky et al., 2010).
Metabolic iron oxydation
Despite the absence of direct evidence, it is becoming increasingly accepted that anoxygenic phototrophic bacteria (photoferrotrophy) were directly involved in the primary oxidation of Fe(II) to Fe(III) in BIF (Konhauser et al., 2002; Kappler et al., 2005;
Posth et al., 2008). This mechanism is strengthen by the identification of modern analogue marine and freshwaters anoxygenic Fe(II)-oxidizing phototrophs (Crowe et al., 2008; Walter et al., 2009). In the past 20 years, a number of experimental studies have confirmed that various purple and green phototrophic bacteria can use Fe(II) as electron donor for CO2 fixation (Widdel et al., 1993; Heising et al., 1999; Straub et al., 1999):
4Fe2+ +CO2 + 11H2O + hy CH2O + 4Fe(OH)3 + 8H+
More recently, it has been demonstrated via experiments and calculations that these organisms would have been capable of oxidizing enough Fe(II) to account for the sedimentary ferric iron flux required to produce large BIF deposition (Konhauser et al., 2002; Kappler et al., 2005).
Critical to this hypothesis is the absence of unequivocal evidence for the existence of Fe(II)-oxidizing phototrophs during Archaean.
UV photooxidation of Fe(II)
UV-induced ferrous oxidation has also been advanced as an important mechanism for the oxidation of soluble Fe(II) (Cairns-Smith, 1978; Braterman et al., 1983):
2Fe2+ + 6H2O + hy → 2Fe(OH)3 + 2H2 + 4H+
Although Fe(II) can be oxidized photochemically in simple aqueous systems (Braterman et al., 1983), the efficiency of UV-dependent oxidation in more complex environments such as seawater has been questioned (e.g., Konhauser et al., 2007).
1.4.3.2 Mechanisms of silica precipitation
Proposed deposition mechanisms include (1) direct precipitation of amorphous silica from seawater or hydrothermal fluids, (2) the precipitation of colloidal hydrous silicates (nontronite), (3) Si adsorption onto Fe-oxyhydroxides.
In the absence of silica-secreting organisms, the Precambrian ocean may have been close to saturation with respect to amorphous silica (Siever, 1992). It then has often been suggested that an amorphous silica gel precipitated as a result of evaporative concentration (e.g., Trendall and Blockley, 1970; Garrels, 1987; Maliva et al., 2005).
Considering a hydrothermal Si source, precipitation is trigger by the cooling of the fluid in contact with seawater. In these cases, chert precursor is of primary origin and its
recrystallisation into microquartz through dissolution-precipitation processes occurred during diagenesis (Knauth, 1994; Marin et al., 2010).
Alternatively, Si reacts with Fe-hydroxides to form hydrous Al-poor Fe-rich silicates (Lascelles, 2007) or adsorbs onto Fe-oxyhydroxides forming a siliceous ferric oxyhydroxide precursor (e.g., Konhauser et al., 2005; Fischer and Knoll, 2009). The cherts of BIF should then be considered as diagenetic in origin.
Then, both the parental water mass (seawater or hydrothermal fluids) and the primary or diagenetic origin of chert layers of BIFs remain debated (see section 1.6.4).
1.4.3.3 BIF deposition mechanisms
As debates on the oxidation processes, the origin of silica and the primary precursor(s) are still open, the deposition and banding mechanisms also remain discussed. Main hypothesis are summarized below and schematised in Table 1.1.
References O2 cyanobacteria photoferrotrophs oversaturation hydrothermal seawater primary diagenetic
x x x 1-5
x x x 6-7
x x x x 8
x x x 9
x x x 10
Chert-layers origin Si source
Fe-oxidation processes
Table 1.1 Schematisation of different BIF deposition models depending on the Fe-oxidation process, Si source and the primary of diagenetic origin of chert layers. Beukes et al., 1990 (1); (2) Klein and Beukes,
1989; (3) Isley (1995); (4)Hamade et al. (2003); (5) Steinhoefel et al. (2009); (6) Lascelles (2007); (7) Krapez et al. (2003); (8) Wang et al. (2009); (9) Fischer and Knoll (2009); (10) Posth et al. (2008).
Two interacting water masses
The most well-known model invokes the interaction of two water masses. The formation of Fe oxides layers is interpreted to reflect periods of intensive upwelling of Fe(II)-rich deep reduced waters or hydrothermal plumes into the photic zone of near-coastal waters where cyanobacteria produced oxygen, which in turn led to the precipitation of a ferric oxyhydroxide precursor (e.g., Klein and Beukes, 1989; Beukes et al., 1990; Isley, 1995;
Hamade et al., 2003; Beukes and Gutzmer, 2008; Steinhoefel et al., 2009) (Fig. 1.9).
Different Fe mineralogies (oxides, silicates, carbonates and sulfides) are then interpreted to reflect a redox gradient in the deposition area (Holland, 2005). Long-distance transport
of hydrothermal solutes during oceanic anoxia (Isley, 1995) may have enabled BIF deposition on continental shelves during high sea level (Simonson and Hassler, 1996).
Banding in iron formation has been argued to reflect alternating chemical precipitation of colloidal iron oxyhydroxides and silica (Garrels, 1987) or continuous evaporative silica precipitation with episodic deposition of iron (Morris, 1993). Chert layers would then record periods of hydrothermal quiescence and Fe-poor sedimentation (e.g., Morris, 1993; Steinhoefel et al., 2009).
Fig. 1.9 Simplified model for the BIF deposition implying two interacting water masses. Upwelling of a reduced Fe-rich hydrothermal plume or bottom seawater into the photic zone where cyanobacteria
produced oxygen inducing the precipitation of Fe-hydroxides.
A common parental water mass and two distinct precursors
Considering BIFs as analogous to modern iron-rich sediments precipitated from deep-sea smokers, Krapež et al. (2003) and Lascelles (2007) infer that both Si and Fe have a hydrothermal origin and that the precursor sediment to BIFs was granular iron-rich hydrothermal muds deposited on the flanks of submarine volcanoes. The model of Lascelles (2007) suggests that hot Fe- and Si-rich hydrothermal fluids are rapidly cooled on contact with cold ocean water, reducing the solubility of elements and producing the precipitation of colloidal particles of Al-poor hydrous iron silicate (nontronite) and iron hydroxides (Fig. 1.10 A). The rapid deposition and abundant included water formed unstable mounds around the vents (Fig. 1.10 B). Slumping of the mounds caused by
compaction, dewatering, gravity sliding and seismic events produced turbidity and density currents (Fig 1.10 C and D).
Fig. 1.10 Deposition of BIF according to the model of Lascelles (2007). A- Cooling of hydrothermal fluids releasing Fe2+ and H4SiO4 into ocean induced precipitation of hydrous iron silicate and iron hydroxides around vent. B-Slumping of the mounds creates turbidity and density currents C and D-deposition from
turbidity and density currents.
From Lascelles, 2007
Deposition from the density currents formed typical finely laminated deposits that subsequently underwent diagenesis. During diagenesis, hydrous iron silicates tend to dissociate into iron oxides and colloidal silica giving rise to the bands of chert (Fig. 1.11).
The cherts of BIF are then considered as diagenetic in origin and developped during burial.
Fig. 1.11 Diagenesis of BIF considering hydrous iron silicates and Fe-hydroxides as precursors. A- Deposition of iron silicates with continuous and discontinuous laminas of iron oxides; B-dissociation of iron silicate into iron oxides and colloidal silica; C-differentiation of mesobands by settling of iron oxides
and accumulation of silica below oxide layers. From Lascelles, 2007
Posth et al. (2008) proposed a radically different model for BIFs deposition despite convergences of a common parental water mass for both Si and Fe and a precipitation of these elements via their own precursors. Invoking photoferrotrophs as responsible for oxidation of the Fe in seawater, they demonstrated that the rate of ferric hydroxides formation by iron-oxidizing microbes is temperature-dependent and decoupled from silica. They showed that phototrophic Fe oxidation rate has an optimal temperature range. At temperature above or below this range, precipitation of iron hydroxides would slower (or cease) while the abiotic silica precipitation induced by silica saturation would continue (or increase in case of temperature drop). Then, natural fluctuations in the temperature of the oceanic photic zone were proposed to have led to the primary layering.
A common parental water mass and a common precursor
Alternatively to direct Si precipitation from seawater, Fischer and Knoll (2009) proposed a model where silicic acid adsorbed onto Fe-hydroxide (Fig. 1.12). They consider the Fe oxydation by anoxygenic photosynthesis where the ferric iron produced would rapidly undergo hydrolysis and precipitate as ferric hydroxides. Dissolved silica readily adsorbs to the ferric hydroxides surface, generating siliceous ferric hydroxides precursor that sank to the sea floor along with organic matter. In sediments, bacterial oxidation of organic matter would induce Fe reduction thus liberating silica. This silica would concentrate into pore fluids and ultimately precipitated and transformed into early diagenetic chert.
Fig. 1.12 Schematic model of BIF deposition according to the model of Fischer and Knoll (2009)
