magnetic mineral diagenesisocean.tongji.edu.cn/space/zhifei/files/2017/12/... · 1. introduction...

Invited review

Magnetic mineral diagenesis

Andrew P. RobertsResearch School of Earth Sciences, The Australian National University, Canberra, ACT 0200, Australia

a b s t r a c ta r t i c l e i n f o

Article history:Received 23 October 2014Received in revised form 21 September 2015Accepted 22 September 2015Available online 3 October 2015

Keywords:DiagenesisMagnetic mineralsOrganic carbonOxidationNitrate reductionIron reductionSulphate reductionMethanogenesisAnaerobic oxidation of methaneRemagnetisation

Reduction–oxidation (redox) reactions occur during burial because sediments contain reactive mixtures ofoxidised and reduced components. Diagenetic chemical reactions represent the approach of all sedimentarycomponents toward equilibrium, and control the long-term stability of sedimentary iron-bearingminerals. Mag-netic minerals are sensitive indicators of sedimentary redox conditions and of changes in these conditionsthrough time, with diagenetic effects ranging from subtle to pervasive. Despite the importance of magnetic min-eral diagenesis in paleomagnetism, rock magnetism, and environmental magnetism, and the usefulness of thesesubjects in the Earth and environmental sciences, there is no systematic single published treatment of magneticmineral diagenesis. This paper is an attempt to provide such a treatment for the full range of diagenetic environ-ments. Magneticmineral diagenesis during early burial is driven largely by chemical changes associated with or-ganic matter degradation in a succession of environments that range from oxic to nitrogenous to manganiferousto ferruginous to sulphidic to methanic, where the free energy yielded by different oxidants decreases progres-sively in each environment. In oxic environments, the most important diagenetic processes involve surface oxi-dation of detrital minerals, and precipitation of Fe3+-bearing minerals from solution. In ferruginousenvironments, the most reactive detrital and authigenic iron oxides undergo dissolution, often mediated by dis-similatory iron-reducing bacteria, which releases Fe2+ that becomes available for other reactions. The Fe2+ in so-lution can diffuse upward where it is oxidised to form new authigenic iron (oxyhydr-)oxide minerals or it canbecome bioavailable to enable magnetotactic bacteria to biomineralise magnetite, generally at the base of theoverlying nitrogenous zone. Alternatively, dissimilatory iron-reducing bacteria can produce extracellularmagne-tite within ferruginous environments. In sulphidic environments, iron-bearing detrital mineral assemblages un-dergo more radical alteration. Hydrogen sulphide, which is a byproduct of bacterial sulphate reduction or ofanaerobic oxidation of methane, reacts with the Fe2+ released from iron mineral dissolution or directly withsolid iron (oxyhydr-)oxide minerals to form iron sulphide minerals (mackinawite, greigite, and pyrite).Authigenic growth of ferrimagnetic greigite has important implications for paleomagnetic recording. Secondaryiron sulphide formation can also occur as a result of anaerobic oxidation ofmethane. Methanemigration throughsediments in association with biogenic or thermogenic methane production or in association with gas hydratedissociation can disrupt the diagenetic steady state and give rise to greigite andmonoclinic pyrrhotite formationthat remagnetises sediments.Most of the above-described diagenetic processes occur below50 °C.With continu-ing burial above 50 °C, but at sub-metamorphic temperatures, magneticminerals can undergo further thermally-induced chemical changes that give rise to awide range ofmineralogical transformations that affect themagneticrecord of the host sediment. These changes include remagnetisations. Magnetic analysis can provide much valu-able information concerning diagenesis in environmental processes. The range of processes discussed in thispaper should assist researchers in analysing sedimentmagnetic properties for which the assessment of diagenet-ic effects has become a necessary component.

© 2015 Elsevier B.V. All rights reserved.

Contents

1. Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22. Nomenclature . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3

2.1. Early versus later diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32.2. Diagenetic zonations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32.3. Iron diagenesis and sediment colour . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5

Earth-Science Reviews 151 (2015) 1–47

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2.4. Steady state and non-steady state diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53. Energetics of organic carbon degradation as a driver of early diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54. Chemical processes that affect magnetic mineral diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5

4.1. Reactivity of iron-bearing minerals . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54.2. Magnetic mineral diagenesis in different diagenetic zones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7

4.2.1. Oxidative magnetic mineral diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 74.2.2. Nitrogenous magnetic mineral diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 104.2.3. Ferruginous magnetic mineral diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 114.2.4. Sulphidic magnetic mineral diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 134.2.5. Methanic magnetic mineral diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19

4.3. Magnetic mineral diagenesis in different sedimentary environments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 214.3.1. Near-shore marine environments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 214.3.2. Hemi-pelagic marine environments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 224.3.3. Pelagic marine environments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 234.3.4. Lake sediments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 254.3.5. Soils . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27

4.4. Non-steady state diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 284.5. Diagenetic microenvironments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 314.6. Silica diagenesis and magnetic minerals . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 324.7. Relict magnetic mineral assemblages in sulphidic sediments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33

5. Physical processes and magnetic mineral diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 355.1. Post-depositional magnetisation lock-in . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 355.2. Sediment compaction and inclination flattening . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 355.3. Locking of sedimentary magnetic fabrics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 36

6. Temperature-dependent diagenetic changes during burial . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 367. Remagnetisation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 378. Outstanding questions concerning magnetic mineral diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 389. Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 38Acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 38References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 38

1. Introduction

Diagenesis refers to the physical and chemical changes that occurduring conversion of sediment into sedimentary rock. The prefix dia-and root -genesis are both from Greek and mean “through” or “across”and the “origin or mode of formation”, respectively. Together, theyrefer to the formation of a sedimentary rock through its entire volume.Sediments undergo considerable change during early burial as theytransform into sedimentary rock. Such so-called “early diagenesis” hasbeen widely studied and relates to processes that occur just below thesediment–water or sediment–atmosphere interface to burial depths ofup to a few hundred metres (e.g., Berner, 1980). Early diagenesis in-cludes chemical and microbial processes, such as reactions associatedwith degradation of organicmatter that involve dissolved chemical spe-cies in pore fluids, and mineral dissolution, precipitation, and cementa-tion, and physical processes, such asmixing by bioturbation, diffusion oradvection of pore fluids, compaction, and dewatering. Most early diage-netic processes are (micro-)biologicallymediated to such an extent thatbiological activity is fundamental to the rate of diagenetic reactions.Sedimentary rocks can also undergomajor physical or chemical changesduring deeper burial, including reactions associated with wide-rangingfluids (meteoric to hydrocarbons), elevated temperatures (above 50°C), and the mechanical consequences of burial at greater depth/pres-sure, with later diagenesis usually taken to occur above 50 °C beforethe onset of metamorphism. Our focus is mainly on diagenetic changesthat occur in the upper several hundredmetres of the sediment column,which are traditionally classified as “early diagenesis”, although there isoften no clean separation between early and later diagenesis for manymagnetically important reactions.

Iron makes up more than 30% of the Earth's mass and is, therefore,one of the most abundant elements on Earth. Iron generally occurs inonly two oxidation states at Earth's surface: as reduced (ferrous) Fe2+

or oxidised (ferric) Fe3+. The terrestrial iron cycle involves a myriad ofreactions where these species can occur in solid form as minerals andin fluids as dissolved ions. Diagenetic reactions involving iron-bearing

fluids and minerals are fundamental to the global iron cycle(e.g., Pérez-Guzmán et al., 2010; Raiswell and Canfield, 2012). Dissolu-tion and precipitation of different minerals are controlled by oxygencontent, reduction potential (Eh), hydrogen ion concentration (pH),and (micro-)biological activity. In oxic aqueous or moist environments,Fe2+ will oxidise rapidly to Fe3+ with a half-life of a few minutes(Stumm and Morgan, 1996). Fe3+-bearing minerals are, therefore,dominant in oxic environments. In contrast, Fe2+-bearing mineralswill only be stable in anoxic environments (at neutral pH). Thus, aniron-bearingmineral will only be in equilibrium (i.e., chemically stable)under a specific range of Eh-pH conditions (Garrels and Christ, 1965;Pourbaix, 1974; Brookins, 1988). Progressive Eh–pH changes duringdiagenesis mean that different chemical stability fields exist for differ-ent iron-bearing minerals in each diagenetic environment (Berner,1980). It is the variable stability of iron-bearingminerals during diagen-esis that is the subject of this paper.

Magnetic minerals are an important class of iron-bearing mineralsthat occur as iron oxides (magnetite, maghemite, haematite),oxyhydroxides (goethite, ferrihydrite, lepidocrocite), and sulphides(greigite, pyrrhotite). These minerals are highly sensitive to diageneticoxidation and reduction reactions, which involve migration of Fe ionsfrom one crystal lattice (e.g., lepidocrocite) to another (e.g., magnetite),and are potentially reversible if environmental conditions change (Liuet al., 2012a). Each iron atom in sediments is estimated to undergo asmany as 100 cycles of reduction and oxidation prior to permanent burial(Thamdrup, 2000). Electron transfer between Fe2+ and Fe3+ is energet-ically favourable (only ~0.01 eV), with each transfer causing a change inmagnetic moment of 1 Bohr magneton (9.27 × 10−24 Am2), which rep-resents amagnetisation change of 25% (Liu et al., 2012a). Magneticmea-surements are used routinely to detect different magnetic minerals.Magnetisation changes associated with iron migration between iron-bearing minerals are, therefore, useful for understanding redox-relatedenvironmental changes.

Magnetic minerals are important for deciphering components of theglobal iron cycle, but they are also fundamentally important for

2 A.P. Roberts / Earth-Science Reviews 151 (2015) 1–47

recording sedimentary paleomagnetic signals that are used widely inthe Earth sciences. While the importance of magnetic mineral diagene-sis, including remagnetisation of sedimentary rocks, has long beenrecognised (e.g., Creer, 1962; Henshaw and Merrill, 1980; Channellet al., 1982; Lowrie and Heller, 1982; Karlin and Levi, 1983), much sed-imentary paleomagnetism has effectively involved magnetic “remotesensing”whereby bulkmeasurements are used to interpret geomagnet-ic and geological phenomena. This approach has involved little mean-ingful interrogation of the nature and microstructure of the magneticminerals that record the measured paleomagnetic signals and thepost-depositional changes to which these minerals may have been sub-jected. As defined above, diagenesis acts through the entire volume ofsediment. It is now widely appreciated that magnetic minerals are notexempt fromdiagenetic alteration. Such effects are ubiquitous, althoughthey vary enormously in scope and extent. It is the ubiquity of diagene-sis that makes it essential to understand and constrain its influence onpaleomagnetic recording. Understanding magnetic mineral diagenesisis fundamental to any field in which magnetic minerals are relevant,particularly in environmental magnetism where magnetic mineralsare used to provide evidence for awide range of environmental process-es (e.g., Thompson and Oldfield, 1986; Verosub and Roberts, 1995;Maher and Thompson, 1999; Evans and Heller, 2003; Liu et al.,2012a). Diagenetic effects can have a dominant control on sedimentmagnetic properties or they can be much more subtle. The author's ex-perience is that diagenesis is almost always an important factor whenanalysing themagnetic properties of sediments to study geochronology,tectonics, geomagnetic field behaviour, or environmental processes. Al-though there is an extensive literature on different aspects of magneticmineral diagenesis, there is no systematic single treatment of the rangeof diagenetic processes that affect magnetic minerals. The purpose ofthis paper is to provide such an overview of magnetic mineral diagene-sis to assist researchers in untangling diagenetic effects that range fromdominant to subtle.

2. Nomenclature

The literature on diagenesis contains variable definitions and classi-fications of processes. This makes it important to define some of themost widely used concepts, as outlined below.

2.1. Early versus later diagenesis

The distinction between early and late diagenesis usually follows thedivide that often exists between “soft” and “hard” rock geologists.Whilethis distinction can be convenient because the processes that affect softsediments and lithified sedimentary rocks are often markedly different,it is not as useful for considering magnetic mineral diagenesis becausethe biological, chemical and physical processes that control magneticmineral diagenesis are often not so readily separated into these simplecategories. Many important processes occur commonly during “early”diagenesis, but some reactions can also occur during “later” diagenesisif the necessary chemical reactants are present. Additionally, whenstudying sedimentary rocks, it is often difficult to determine whetherauthigenic mineral phases grew during early or later diagenesis. Lateformation of magnetic minerals gives rise to remagnetisations(e.g., Jackson et al., 1988; McCabe and Elmore, 1989; Jackson andSwanson-Hysell, 2012; van der Voo andTorsvik, 2012),which in ancientrocks are often assumed to have been recorded during late diagenesis.However, microtextures associated with authigenic pyrite and greigitecan be apparently identical for modern sediments (e.g., Rowan et al.,2009), and those that apparently record paleomagnetic signals thatwere acquired within several tens of thousands of years (e.g., Robertset al., 2005; Larrasoaña et al., 2007) or up tomillions of years after depo-sition (e.g., Roberts and Weaver, 2005; Rowan and Roberts, 2006). Inthese cases, remagnetisations are typically detected using paleomagnet-ic field tests such as the fold test (Rowan and Roberts, 2006, 2008) or by

detecting contradictory paleomagnetic polarities (e.g., Jiang et al., 2001;Sagnotti et al., 2005a). The difficulty in conclusively identifying thetiming of authigenic magnetic mineral growth in poorly lithifiedsedimentary rocks has led to attribution of remagnetisations as occur-ring during “later” diagenesis (e.g., Jiang et al., 2001; Roberts andWeaver, 2005; Roberts et al., 2005, 2010; Rowan and Roberts, 2006).This term is vague, but it recognises the reality that the diagenetic mag-netic mineral transformations of interest fall within neither of the tradi-tionally recognised early or late diagenetic stages. Throughout thispaper, the terms early, late and later diagenesis are used in the relativis-tic sense defined here.

2.2. Diagenetic zonations

Post-depositional degradation of organic matter is a fundamentallyimportant process for driving early diagenetic chemical changes in sed-imentary environments. Sedimentary organic matter accumulatesthrough supply of reactive organic matter of various types, includingterrestrial and/or marine plant, animal and planktonic remains. Micro-bial respiration of organic matter in the water column and sediments,wheremicrobes derive energy through oxidation of organic compoundsto release CO2 from the organicmatter, leads to oxidative degradation ofreactive organic compounds in the organic matter. Microbial metabo-lism of sedimentary organic matter proceeds through a sequence inwhich the free energy yielded by different oxidants (permole of organiccarbon oxidised) progressively decreases (Froelich et al., 1979)(Table 1). The order of electron acceptor use is as follows: oxygen, ni-trate, manganese oxides, iron (oxyhydr-)oxides, sulphate, and organicmatter itself. When one oxidant is depleted, the next most efficient(i.e., most energy producing) oxidant is used, etc., until either all oxi-dants or all reactive organic matter are consumed (Froelich et al.,1979). Progressive consumption of these electron acceptors, and forma-tion of the products of the reactions (Mn2+, Fe2+, H2S, CH4), are illus-trated in an idealised diagram in Fig. 1 and the free energiesassociated with electron transfer for the respective reactions are listedin Table 1. That is, as manganese oxides, iron (oxyhydr-)oxides, and sul-phate are reduced, the concentration of Mn2+, Fe2+, and H2S, respec-tively, increases in sedimentary pore waters. Once pore watersulphate is consumed entirely via microbial sulphate reduction and byanaerobic oxidation of methane, the dominant process by which organ-ic matter is degraded is via methanogenesis (Fig. 1). Diagenetic libera-tion of dissolved Mn2+, Fe2+, H2S, and CH4 into sedimentary porewaters provides reactants for authigenic mineral formation; the

Table 1Gibbs free energy for respiratory pathways of organic matter remineralisation.1,2

Respiratory pathway with equations kJ/reactionΔG (acetate)3

Oxic respiration −402O2+0.5C2H3O2

−→0.5H++HCO3−

Denitrification −3590.6H++0.8NO3

−+0.5C2H3O2−→0.4N2+HCO3

−+0.2H2OMn reduction4 −3853.5H++2MnO2+0.5C2H3O2

−→2Mn2++HCO3−+2H2O

Fe reduction −2417.5H++4FeOOH+0.5C2H3O2

−→4Fe2++HCO3−+6H2O

Sulphate reduction −43.80.5H++0.5SO4

2−+0.5C2H3O2−→0.5H2S+HCO3

−

Methanogenesis −19.90.5H2O+0.5C2H3O2

−→0.5CH4+0.5HCO3−

1. After Canfield and Thamdrup (2009).2. Acetate is used as the electron donor.3. Calculated values are standardised to a four-electron transfer equivalent to oxidation of1 mol of organic carbon as a carbohydrate (at 25 °C, pH= 7, and unit activity for all reac-tants and products).4. The value calculated for Mn reduction is higher than for denitrification; the exact valuewill depend on the mineral phase presumed to be undergoing dissolution and can varyfrom slightly higher to slightly lower values than for denitrification (e.g., Burdige, 1993).

3A.P. Roberts / Earth-Science Reviews 151 (2015) 1–47

presence of authigenic minerals in the geological record is, therefore, anindicator of former early diagenetic conditions (Berner, 1981).Authigenic formation of the most relevant iron-bearing minerals in dif-ferent diagenetic zones is illustrated on the right-hand side of Fig. 1.

The diagenetic zones produced by the above-described reactionshave been traditionally referred to as follows: (1) oxic—the upperzone in which organic matter respiration occurs under oxic conditionswhere measurable concentrations of oxygen are present and wherethe terminal electron acceptor associated with organic matter degrada-tion is oxygen; (2) suboxic—the intermediate zone inwhich neither ox-ygen nor dissolved sulphide are present in measurable concentrationsand in which nitrate, manganese, and iron are successively reduced;and (3) anoxic—the lowermost zone where oxygen is strictly excludedand that is often subdivided into an upper sulphidic zone and a lowermethanic zone in which pore water sulphide concentrations havebeen reduced to zero (Froelich et al., 1979; Berner, 1981). In some set-tings, the sulphidic andmethanic zones can overlap significantly, partic-ularly those with low reactive iron oxide inputs (Niewöhner et al.,1998). This nomenclature concerning diagenetic zones has been usedwidely for decades and is embedded in the literature—I have used thisset of definitions throughout my own work on magnetic mineral dia-genesis (most recently in relation to diagenesis in pelagic marine car-bonates; Roberts et al., 2013a). Canfield and Thamdrup (2009) arguedthat this nomenclature is inconsistent and confusing for a range ofreasons, particularly in relation to the term suboxic. For example, insome continental margin settings or deeper water sulphidic basinswith high organic carbon contents the redox gradient is so steep thatoxygen is depleted belowdetection limits in themanganese and iron re-duction zones where pore waters are anoxic rather than ‘suboxic’(e.g., Christensen et al., 1989; Murray et al., 1989; Canfield et al., 1993;Kuypers et al., 2003; Percy et al., 2008). In other cases, manganese re-duction (e.g., Canfield et al., 1995) or sulphate reduction (Canfield andDes Marais, 1991; Jørgensen and Bak, 1991) have been documented infully oxygenated waters. Likewise, definition of the term anoxic isbased on the lowest O2 concentration measurable with standard tech-niques, but modern sensors canmeasure O2 concentrations with orders

of magnitude greater sensitivity than was the case 30 years ago. Thismakes it difficult to compare definitions between older and more mod-ern studies. Examples of other such inconsistencies are provided byCanfield and Thamdrup (2009), who proposed abandonment of theterms suboxic and anoxic as defined by Froelich et al. (1979).

Canfield and Thamdrup (2009) proposed a diagenetic zonationbased on well-recognised chemical environments and associated mi-crobial processes so that emphasis is on specific metabolic processesrather than on non-specific, imprecise, and inconsistently appliedterms like ‘suboxic’ metabolism. In this paper, the scheme of Canfieldand Thamdrup (2009) is used, where, from top to bottom, the chemicalzones and accompanying respiration processes (in parentheses) are(Fig. 1): oxic (aerobic respiration), nitrogenous (nitrate reduction),manganous (manganese reduction), ferruginous (iron reduction),sulphidic (sulphate reduction), and methanic (methanogenesis).These zones are defined by well-known processes that can be relatedeasily to published studies, but with more precise language.

Iron-bearingminerals are intimately associated with organic matterdecomposition irrespective of the type of sedimentary setting (Fig. 1,right-hand side). In ferruginous environments, dissolved iron concen-trations increase in sedimentary pore waters. If the dissolved iron dif-fuses upward into nitrogenous or aerobic environments, it can beoxidised to cause precipitation of iron oxides (magnetite (Fe3O4), hae-matite (α-Fe2O3)), oxyhydroxides (goethite (α-FeOOH), lepidocrocite(γ-FeOOH), ferrihydrite (5Fe2O3 • 9H2O)), and transient iron hydroxylsalts known as green rusts. In sulphidic environments, sulphate general-ly decreases to near-zero values and dissolved sulphide (H2S) concen-trations increase. H2S reacts with dissolved iron to form paramagneticpyrite (FeS2), including precursor phases such as mackinawite(Fe1+xS) and ferrimagnetic greigite (Fe3S4). Inmethanic environments,dissolved iron can build up in interstitial waters, but it cannot react toform pyrite because dissolved sulphide is no longer produced. The re-duced iron is, therefore, available to precipitate with other dissolvedions. If interstitial waters are saturated with respect to carbonate, para-magnetic siderite (FeCO3) can form. In this geochemical scheme,authigenic iron-bearing minerals can be diagnostic of the diagenetic

Fig. 1. Cartoon representation of the depth distribution of sedimentary redox-driven diagenetic zones. Electron acceptors and respiration processes by which reactants are consumed areindicated on the left. Idealised porewater profiles of reactants (O2, NO2

‐ , NO3‐) and products (NO3

‐ , Mn2+, Fe2+, H2S, CH4) and associated chemical zones are shown on the right (modifiedfrom Canfield and Thamdrup (2009) and Jørgensen and Kasten (2006)). The names from Canfield and Thamdrup (2009) for chemical zones are used, with the same colour coding,throughout this paper. Authigenic iron minerals that can form in the respective chemical zones are listed in the far right-hand column (from Berner, 1981).


stages through which the sediment progressed (Berner, 1981) and canbe used to interpret former diagenetic conditions in ancient sediments.

2.3. Iron diagenesis and sediment colour

Sediment colour is controlled by several variables, including organiccarbon content, the oxidation state of iron and the colour of the mainconstituent rock-forming minerals (Giosan et al., 2002; Stow, 2005).Oxic sediments are typically white, yellow, tan, brown or red. Inmarinesediments, a colour transition accompanies the oxic/nitrogenous to fer-ruginous diagenetic transition, with brown colours giving way to olivegreens in association with reduction of Fe3+ to Fe2+ (Lyle, 1983;Giosan et al., 2002). In sulphidic environments, reactive Fe3+ is reducedto Fe2+ and sediments have a grey appearance. With increasing organiccarbon content, colours range from light to dark grey to black; dark col-ours can also result from the presence of finely dispersed pyrite andmonosulphides (Stow, 2005). Colour variations are, therefore, a usefulindicator of mineral components and diagenetic alteration. In the fig-ures below, oxic sediments are indicated with brown, ferruginous sedi-ments are indicated with olive green, and sulphidic sediments areindicated with grey.

2.4. Steady state and non-steady state diagenesis

Most treatments of early diagenesis, including that presented above,involve the assumption that diagenesis proceeds under steady-stateconditions (e.g., Berner, 1980). That is, steady state diagenesis occurswhen sediment input, including inorganic sediment and reactive organ-ic matter, remains approximately uniform so that progressive diagenet-ic changes with depth are attributed solely to time-dependent chemicalprocesses (Raiswell and Canfield, 2012). This idealised conception ishelpful for considering many diagenetic scenarios. In contrast, non-steady state diagenesis occurs when sediment input varies with timeso that diagenetic reactions proceed in one direction and then changeto another as depositional boundary conditions change (e.g., Bergeret al., 1983; Thomson et al., 1984). Variable organic carbon supply tothe seafloor through orbitally controlled variation in surface ocean pro-ductivity (e.g., Dekkers et al., 1994; van Os et al., 1994; van Santvoortet al., 1997; Larrasoaña et al., 2003a), or through instantaneous deliveryof elevated organic carbon contents from the continental shelf to thedeep-sea by turbidites (Robinson and Sahota, 2000; Robinson et al.,2000), provide good examples of how diagenetic environments cancycle between oxic and sulphidic conditions to produce non-steadystate diagenesis. In this paper, steady-state diagenesis is assumed unlessnon-steady state diagenesis is mentioned specifically. In reality, non-steady state diagenesis is often more important than steady state dia-genesis because organic carbon delivery or sedimentation rate are notconstant. This should be borne in mind when considering diageneticscenarios.

3. Energetics of organic carbon degradation as a driver ofearly diagenesis

As outlined above, degradation of reduced carbon in organic mattergenerally proceeds through microbial exploitation of the energy gainedby using dissolved and solid phase oxidants (oxygen, nitrate, iron andmanganese oxides, and sulphate) as terminal electron acceptors(Raiswell and Canfield, 2012). The well-known order of terminal elec-tron acceptor use (Table 1) is based on thermodynamic calculations ofthe free energy yielded (e.g., Froelich et al., 1979; Stumm and Morgan,1996; Canfield and Thamdrup, 2009); pore water data support the as-sertion that electron acceptors are used in the order of decreasing freeenergy yield (e.g., Froelich et al., 1979; Martin and Sayles, 2005). As in-dicated in Table 1, the difference in free energy yielded by denitrifica-tion and manganese reduction is not large; whether one processyields more energy than the other will depend on the manganese-

bearing mineral phase undergoing dissolution (e.g., Burdige, 1993), sothat the order of the two processes can vary (Martin and Sayles,2005). Overall, however, when one oxidant is consumed, the next isused, and then the next, etc., until the diagenetic system approachesequilibriumwhen all reactive organic carbon is oxidised. Organicmatteroxidation occurs at different rates and to different extents in sedimenta-ry environments with contrasting organic carbon supply. For example,in pelagic marine environments, organic matter supply tends to below and respiration tends to occur dominantly through aerobic oxida-tion (N90%), followed by nitrate reduction and manganese reductionwithout progressing further (Table 2). In contrast, continental marginsediments tend to have higher organic carbon contents so that iron re-duction, sulphate reduction and methanogenesis are also importantprocesses (Table 2). This variability is responsible for the contrasting in-fluence of diagenesis onmagnetic minerals in different depositional en-vironments, as discussed below.

4. Chemical processes that affect magnetic mineral diagenesis

Much of the treatment of magnetic mineral diagenesis in this paperfocuses on the effects of chemical change during diagenesis (Section 4).This treatment begins with consideration of the reactivity of iron-bearingminerals (Section 4.1), followed by discussion of magneticmin-eral diagenesis in the respective diagenetic zones (Section 4.2). The ex-tent to which pore waters evolve to reveal the entire set of processesillustrated in Fig. 1 depends on the rate of organic carbon supply andhow quickly it can be oxidised before each oxidant is progressively con-sumed. This diagenetic progression usually varies significantly betweenenvironments (Table 2), so it is also useful to discuss diagenesis in termsof depositional environment (Section 4.3). For example, freshwater typ-ically has dissolved sulphate contents that are b1% of that of seawater(Capone and Kiene, 1988). Reductive diagenesis, therefore, often fol-lows different pathways in freshwater compared to marine sediments.The discussion throughout Sections 4.1-4.3 revolves largely aroundsteady state diagenesis. Non-steady state diagenesis is considered ex-plicitly in Section 4.4. Diagenetic microenvironments are considered inSection 4.5, silica diagenesis is considered in relation to magnetic min-erals in Section 4.6, and relict magnetic mineral assemblages insulphidic sediments are discussed in Section 4.7. Physical processesthat are strictly related to early diagenesis, such as sediment compactionand remanent magnetisation acquisition, are discussed in Section 5,while temperature-dependent diagenetic changes during burial are de-scribed in Section 6. Remagnetisations are considered in Section 7.

4.1. Reactivity of iron-bearing minerals

Redox reactions occur during early diagenesis because sedimentscontain reactive mixtures of oxidised and reduced components; diage-netic reactions represent the approach of the total assemblage towardequilibrium (e.g., Raiswell and Canfield, 2012). Iron-bearing mineralsare variably reactive, whichmakes it important to understand their dif-ferential reactivity when consideringmagnetic mineral diagenesis. “Re-active iron” was initially defined by Berner (1970) as the iron releasedby extraction during a 1-min boiling in concentrated HCl. Canfield(1989) used this definition and a dithionate extraction to link the reac-tivity of a range of synthetic iron oxides to dissolved sulphide and iden-tified a highly reactive pool of ferrihydrite and lepidocrocite and a lessreactive pool of goethite and haematite. Crucially, dissolved sulphideis absent from pore waters when both pools of iron oxides are present,but once much of the exposed reactive iron oxyhydroxide pool is con-sumed, dissolved sulphide is present. This means that ferrihydrite andlepidocrocite are likely to bemost reactive in ferruginous environmentsand that goethite and haematite will start to react in sulphidic environ-ments. In sulphidic environments, all of these minerals react rapidlywith dissolved sulphide to form iron sulphides (Canfield, 1989).Canfield et al. (1992) observed that long-term exposure to pore water


sulphide causes reaction of iron-bearing sheet silicate minerals (biotite,chlorite, smectite, illite) at rates 108 times as slow as for the most reac-tive iron oxyhydroxides. This differential reactivity of iron-bearingmin-erals to sulphide explains why pore water sulphide can accumulate insediments (Fig. 1) after iron oxides have been dissolved to form pyrite.Canfield et al. (1992) determined a reaction series of iron-bearing min-erals to a 1 mMdissolved sulphide concentration as follows (frommostto least reactive): ferrihydrite→ lepidocrocite→ goethite→ haematite→ magnetite → “reactive” silicates → sheet silicates →ilmenite, garnet, augite, amphibole. This reaction series is important be-cause it explains why magnetic minerals like haematite and magnetiteare rare (e.g., Karlin and Levi, 1983, 1985; Canfield and Berner, 1987)and why ilmenite is common in diagenetically reduced sediments(e.g., Roberts and Turner, 1993; Wilson and Roberts, 1999; Nowaczyk,2011).

In contrast to Canfield et al. (1992), haematite and/or goethite havebeen reported to bemore stable to reductive dissolution thanmagnetitein sedimentary magnetic studies (e.g., Yamazaki et al., 2003; Liu et al.,2004; Emiroglu et al., 2004; Rey et al., 2005; Kawamura et al., 2007;Rowan et al., 2009; Mohamed et al., 2011). This probably results fromthe fact that dissolution reaction kinetics are difficult to determine be-cause iron minerals can have variable iron concentration, surface area,and differing competitive effects of dissolved species (Poulton et al.,2004). In a revised scheme for the reactivity of iron-bearing mineralsto sulphide, Poulton et al. (2004) reported two groups of mineralswith different reactivity: a more reactive group that includes hydrousferric oxide, lepidocrocite and ferrihydrite, and a less reactive groupthat includes goethite, magnetite and haematite (Table 3). In this

scheme, magnetite is more reactive than haematite, and similarly reac-tive to goethite. Poulton et al. (2004) attributed this higher reactivity ofmagnetite (by a factor of 50) to differences in grain size (and, hence,surface area) in their study compared to the coarser magnetite consid-ered by Canfield et al. (1992). It should also be noted that the relative re-activity of haematite and goethite is reversed in some studies (Postma,1993; Poulton et al., 2004), again due to surface area effects. Overall, re-activities are similar among the two groups of minerals and range fromminutes to hours for one group and tens of days for the less reactivegroup (Table 3). Variable reactivities reported in magnetic studies areconsistent with the variable effects of different parameters cited inthese studies. This should be borne inmindwhen interpretingmagneticresults from sediments. Combining the updated results of Poulton et al.(2004) with those from Canfield et al. (1992) for more refractory min-erals gives a hybrid revised reactivity scheme (from most to least reac-tive; Table 3): hydrous ferric oxide N lepidocrocite ≈ ferrihydrite ≫goethite Nmagnetite N haematite≫ “reactive” silicates≫ sheet silicates≫ ilmenite, garnet, augite, amphibole.

Understanding the reactivity of iron-bearing minerals to dissolvedsulphide is important for several reasons. First, it helps to understandthe sequence of reactions expected and why some minerals are, andothers are not, preserved in particular diagenetic environments. For ex-ample, dissolved iron typically occurs in nanomolar concentrations inthe deep oceans (e.g., Boyd and Ellwood, 2010). Thus, for iron to be bio-available to organisms such as magnetotactic bacteria at sufficient con-centrations to enable magnetite biomineralisation, diagenetic releaseand upward diffusion of iron within the sediment become important.Diagenetic iron release is most likely to occur within ferruginous sedi-ments where iron reduction liberates iron from the most labile iron-bearing minerals (hydrous ferric oxide, lepidocrocite, ferrihydrite) andmagnetite biomineralisation by magnetotactic bacteria is most likelyto occur after the released Fe2+ has diffused up to the base of the over-lying nitrogenous zone. If the inorganic remains ofmagnetotactic bacte-ria (magnetofossils) are preserved within the sediments, this meansthat sulphidic conditions cannot have formed because the magnetiteproduced by the bacteria would have dissolved and would not be pre-served in the geological record (e.g., Roberts et al., 2011a). Second, al-though pyritisation in sediments has often been assumed to be ironlimited (i.e., where rates of dissolved sulphide production exceed therate of iron burial so that all reactive iron is transformed to pyrite),iron limitation is a relative term that depends on the reactivity of iron-bearing minerals with respect to rates of sulphide production(Canfield et al., 1992). Thus, preservation of intermediate magneticiron sulphides like greigite, whichdid not transform completely into py-rite, can be related to an abundance of reactive iron relative tomore lim-ited production of dissolved sulphide (e.g., Kao et al., 2004). Third, slowreaction of more refractory iron-bearing phyllosilicate minerals withpore water sulphide over long periods of time can give rise to slowbut continuous iron sulphide formation, which can have important

Table 2Variation in organic matter oxidation and electron acceptor use in marine sediments.

Site Other information Corg oxidation rate % Corg oxidation by electron acceptor

(μmol cm−2 year−1) O2 NO3− Mn4+ Fe3+ SO4

2−

Pelagic sedimentsMANOP H1 E. eq. Pacific 12.0 99.2 0.8 0.4MANOP C1 Central eq. Pacific 20.4 98.1 1.6 0.4E. Eq. Atlantic1 12.4 93.8 4.4 0.1 1.8

Continental margin sedimentsNE Atlantic2 36–158 67–97 1–8.5 0–2.1 0–1.7 1–20NW Atlantic3 36–52 74–90 1.8–6.0 8–20NE Pacific4 O2 b 50 μM 66–75 5–46 41–69 0.1 0.7–1.3 5.7–25NE Pacific4 O2 = 73–145 μM 36–74 69–75 11–18 0.1–6.9 0.3–0.7 5.6–18

This table contains a selected range of environments to exemplify the variable significance of different electron acceptors. 1. From Bender andHeggie (1984). 2. From Lohse et al. (1998). 3.From Martin and Sayles (2004). 4. From Reimers et al. (1992).

Table 3Half-lives (t1/2) for reductive dissolution of iron-bearing minerals.

Mineral t1/2

Short half livesFresh hydrous ferric oxide 5.0 minLepidocrocite 10.9 h2-line ferrihydrite 12.3 h

Longer half livesGoethite 63 daysMagnetite 72 daysHaematite 182 days

Refractory minerals“Reactive” silicates 333 yearsSheet silicates 122,000 yearsIlmenite, garnets, augite, amphibole NN122,000 years

Data for hydrous ferric oxide to haematite are from Poulton et al. (2004); for seawater atpH= 7.5, 25 °C, and 1 mM dissolved sulphide concentration. Data for “reactive” silicatesto ilmenite, garnets, augite and amphibole are from Canfield et al. (1992); for seawater atpH= 7.5 and 1 mM dissolved sulphide concentration. A 1 mM sulphide concentration isused because this value is typical of porewaters whenmost sulphate has been consumed.


consequences for authigenic magnetic mineral formation and for thepaleomagnetic record of sediments (e.g., Jiang et al., 2001; Robertsand Weaver, 2005).

4.2. Magnetic mineral diagenesis in different diagenetic zones

Diagenetic alteration and/or precipitation of iron-bearing mineralsoccur in each of the oxic, nitrogenous, ferruginous, sulphidic, andmethanic zones (Fig. 1). The following discussion focuses on thesezones, with relatively little emphasis on the manganous zone.

4.2.1. Oxidative magnetic mineral diagenesisDiagenetic changes that affect magnetic minerals in oxidising sedi-

ments principally involve diffusive oxidation ofmineral surfaces bymo-lecular oxygen or precipitation of Fe3+-bearing minerals (Fig. 2) due toupward diffusion of Fe2+ ions that are oxidised at the base of the nitrog-enous zone (Fig. 1). These processes are discussed in separate sub-sections below.

4.2.1.1. Oxidation of mineral surfaces. As amixed valence iron oxide, mag-netite is thermodynamically unstable in oxidising environments. Atambient temperatures, magnetite oxidises to maghemite; at higher tem-peratures, it oxidises to haematite. Partial (surficial) oxidation of magne-tite has been commonly reported in terrestrial eolian deposits such as theChinese loess (e.g., Liu et al., 2007), in oxic deep-sea sediments(e.g., Henshaw and Merrill, 1980; Torii, 1997; Smirnov and Tarduno,2000; Passier and Dekkers, 2002; Yamazaki et al., 2003; Kawamura

et al., 2007, 2012; Chang et al., 2013), and during incipient weatheringof subaerially exposed marine sediments (van Velzen and Zijderveld,1995). Torii (1997) suggested that rapid maghemitisation occurs at thesediment–water interface in marine sediments, whereas Smirnov andTarduno (2000) and Kawamura et al. (2012) favour ongoing low-temperature oxidation during burial. Low-temperature oxidation affectsboth biogenic (e.g., Smirnov and Tarduno, 2000; Chang et al., 2013) anddetrital magnetite. Maghemitisation proceeds from the surface of parti-cles inward, where oxygen diffuses into the particle and Fe2+ migratesout of the spinel lattice to create lattice vacancies. This gives rise to a lat-ticemismatch, where the surficial maghemite has a smaller crystal latticesize than the magnetite core of the particle (Petersen and Vali, 1987;Housden and O'Reilly, 1990). Irregular, curved shrinkage cracks, there-fore, develop on the surface and outer parts of these composite particles(Fig. 3; Petersen and Vali, 1987). The Verwey transition in magnetitealso becomes progressively suppressed with oxidation (Özdemir et al.,1993; Cui et al., 1994; Özdemir and Dunlop, 2010), which can compro-mise the low-temperature test of Moskowitz et al. (1993) for identifyingmagnetite magnetofossils (Smirnov and Tarduno, 2000; Chang et al.,2013). Other low-temperature characteristics of maghemite (Özdemirand Dunlop, 2010) have, therefore, been proposed as an alternative testfor the presence of partially oxidised magnetite magnetofossil chains insediments (Chang et al., 2013). Overall, progressive maghemitisation ofmagnetite has a major effect on sediment magnetic properties both as aresult of transportation from source to sink in oxic environments (Liuet al., 2007) and as a result of oxic diagenesis (Torii, 1997; Smirnov andTarduno, 2000; Kawamura et al., 2012). Surficially maghemitised

Fig. 2. (a) Schematic illustration of precipitation of Fe3+-bearing minerals at an oxidation front flowing through porous sandstone where oxygenated and reduced Fe2+-bearing fluidsmeet. Reduced sandstone to the left will be bleached, while oxidised sandstone to the right will have red-yellow colours. Modified from Beitler et al. (2005). (b) Schematic illustrationof redox reactions in porous sediments bywhich an initial red sediment (left) is progressively bleached by influx of reducing fluids (second and third panels), and is then oxidised (fourthpanel from left) to form localised Fe3+-oxide concretions (right). Modified from Chan et al. (2006).


magnetite particles are, therefore, both a common depositional constitu-ent as well as a diagenetic product in oxic environments.

It is important to testwhether partial or completemagnetite oxidationaffects paleomagnetic recording. High quality magnetostratigraphic re-cords have been obtained (e.g., Florindo and Roberts, 2005) from sedi-ments that contain magnetite magnetofossils, but for which low-temperature tests indicate maghemitisation and complete loss of aVerwey transition (Roberts et al., 2012). High quality paleomagnetic re-cording is generally observed for similar oxic pelagic carbonates(Roberts et al., 2013a). It is unclear whether short-period paleomagneticsignals are compromised by magnetofossil maghemitisation, but theslowly deposited pelagic carbonates studied by Florindo and Roberts(2005) do not provide a useful test of this possibility. A better test isprovided by Arctic Ocean sediments in which partial oxidation gives riseto particles with a titanomagnetite core and titanomaghemite rim(Channell and Xuan, 2009; Xuan and Channell, 2010). This oxidationhas been argued to cause self-reversalwhere the titanomaghemite carriesa chemical remanent magnetisation (CRM) and the host titanomagnetitecarries a detrital remanent magnetisation (DRM). Paleomagnetic record-ing complexities associated with partial oxidation of titanomagnetitehave been proposed as the cause of spurious geomagnetic excursions(Channell and Xuan, 2009; Xuan and Channell, 2010) reported from theArctic Ocean (e.g., Nowaczyk and Antonow, 1997). Magnetic mineral ox-idation, therefore, has apparently variable effects on the paleomagneticrecord of sediments and care is required to assess its possible influence.

An additional type of mineral surface oxidation that is important inpaleomagnetic studies occurs when sediments or sedimentary rocksthat have undergone sulphidic diagenesis are tectonically uplifted,eroded, and exposed in subaerial weathering environments. Exposureof minerals that contain reduced iron, such as pyrite, to molecular oxy-genwithinmeteoric groundwater or the atmosphere can give rise to ox-idation of pyrite to form goethite or haematite. For example, Lowrie andHeller (1982) and Johnson et al. (1984) outlined several cases in ancientlimestones where goethite carries a late diagenetic paleomagnetic sig-nal due to pyrite oxidation. Likewise, pyrite oxidation has been invokedas responsible for widespread high coercivity overprints in siliciclasticsediments from New Zealand (Turner, 2001; Rowan and Roberts,2006), where the unblocking temperatures of the overprint are toohigh to be due to goethite and are attributed to fine-grained haematite.

4.2.1.2. Precipitation of Fe3+-bearing magnetic minerals from solution.Redox fronts, where dissolved species come into contact with chemicalconditions that allow mineral precipitation or dissolution, are

fundamental to diagenesis. Such fronts can move through the sedimentby diffusion or advection. For example, if Fe2+ occurs in solution due toiron reduction, it can diffuse or be advected laterally or vertically to a lo-cationwhere it comes into contactwithmolecular oxygen in the nitrog-enous zone (Fig. 1). This location becomes a three-dimensional redoxfront at which Fe3+-bearing solid phases can precipitate (Fig. 2). Pre-served redox fronts are widely observed (Fig. 4a–c) as yellow-brown-red coloured zones of iron (oxyhydr-)oxide mineralisation in rock out-crops, sediments, and soils.When Fe2+ comes into contact with oxygen,it is oxidised to form nanoparticulate Fe3+ oxyhydroxide and ferrihy-drite aggregates (Cornell and Schwertmann, 1996), which can be con-verted to goethite and eventually to haematite as follows:

2Fe2þ þ 0:5O2 þ 2H2O ¼ Fe2O3 þ 4Hþ: ð1Þ

Nanocrystalline iron (oxyhydr-)oxides have been argued to coarsen inaqueous environments by self-assembly to form larger crystal sizes viarotation so that their lattice structures adopt parallel orientations andaggregate through formation of Fe–O bonds (Banfield et al., 2000). Thestability of these nanoparticulate iron (oxyhydr-)oxides can be size de-pendent, so that coarsening ferrihydrite can transform into goethite(Banfield et al., 2000).

Precipitation of haematite and/or goethite has been widely docu-mented at the base and margins of oxidising zones in oxic diageneticaqueous environments, including pelagic (e.g., Channell et al., 1982)and hemi-pelagic (Bouilloux et al., 2013) marine sediments, and terres-trial aquifers (Lovley et al., 1990; Chan et al., 2000; Beitler et al., 2005).In all cases, dissolved Fe2+ is generated by reduction of iron from detri-tal, biogenic or authigenic iron-bearing minerals. Diagenetic release ofFe2+ from sediments that support active iron reduction plays an impor-tant role in iron cycling in continentalmargins (e.g., Johnson et al., 1999;Bruland et al., 2005; Lam et al., 2006; Scholz et al., 2014).While hydrousferric oxide, ferrihydrite and lepidicrocite are the most reactive iron(oxyhydr-)oxides in reducing environments (Poulton et al., 2004),when the Fe2+ released by reduction of these minerals is re-oxidisedto Fe3+, it can re-precipitate as authigenic hydrous ferric oxide, ferrihy-drite, lepidicrocite, haematite or goethite depending on Eh/pH condi-tions (e.g., Chan et al., 2000; Beitler et al., 2005).

Diagenesis associated with vertically and laterally variable redoxconditions is important for interpreting colour variations in, and the or-igin of, red sediments. Iron-reducing bacteria give rise to dissolution ofburied Fe3+ oxides wherever they are found, including within deepaquifers. Subsequent groundwater migration from such deep aquifersand late diagenetic precipitation of Fe3+ when the fluids encounteroxidising conditions have been argued to give rise to formation of varie-gated red beds (Lovley et al., 1990). Likewise, an inverted redox zona-tion, due to diffusion of oxygen from seawater circulating in anoceanic basement aquifer into overlying pelagic ferruginous sediments,has been interpreted to give rise to a deep iron oxidation front that,coupled with silica diagenesis, is proposed to explain the origin of redcherts (Meister et al., 2014). Silica precipitation is proposed to havebeen catalysed by adsorption to freshly precipitated iron oxide surfaces,so that iron diagenesis provides a key to chert formation. The deepredox front discussed by Meister et al. (2014) occurred within middleMiocene sediments, so later diagenetic reactions that involve ironoxide precipitation are potentially important for paleomagnetic studies.

Red beds are an important class of sediments for paleomagneticstudies. Early paleomagnetic studies of red beds recognised that haema-tite, which gives rise to their red colouring, occurs in both detrital andauthigenic forms. This led to a major controversy in the late 1970s/early 1980s, which became known as the ‘red bed wars’, concerningthe reliability of paleomagnetic records carried by red sediments. Onecamp argued that specular haematite has a detrital origin and that it,therefore, carries a syn-depositional DRM (e.g., Elston and Purucker,1979; Steiner, 1983; Herrero-Bervera and Helsley, 1983). The othercamp argued that red beds have multiple overlapping magnetisation

Fig. 3. Scanning electron microscope (SEM) image (backscattered electrons) of asurficially oxidised detritalmagnetite particle. Maghemitisation proceeds from the surfaceof particles inward,with O2 diffusion into the particle and Fe2+migration out of the spinellattice to create lattice vacancies. The surficial maghemite has a smaller crystal lattice unitcell than the magnetite core of the particle (8.33–8.35 Å versus 8.397 Å, respectively),which gives rise to the observed irregular, curved shrinkage cracks on the outer part ofthe composite particle (Petersen and Vali, 1987). The sample is from a magnetic extractof sediments from Lake Kinneret, Israel (modified from Nowaczyk (2011)).


components carried by chemically precipitated pigmentary haematitethat produces a CRM that could have been acquired over millions ofyears (e.g., Larson and Walker, 1975; Turner, 1980; Larson et al.,1982). Application of modern paleomagnetic analysis, with detailedstepwise demagnetisation and principal component analysis to identifycharacteristic remanent magnetisation components, can be used to re-solve these issues because pigmentary and specular haematite producecontrasting unblocking temperature spectra (van der Voo and Torsvik,2012). Pigmentary haematite typically undergoes gradual unblockingand specular haematite tends to undergo sharper unblocking at theNéel temperature of haematite, 680 °C (Collinson, 1974; van der Vooand Torsvik, 2012). Red beds can record paleomagnetic signals due to

both a DRM and a CRM, and later haematite growth can causeremagnetisation of a primary DRM. Distinguishing between these possi-bilities is a key aspect of paleomagnetic analyses of red sediments.

While haematite pigment is common in sediments and sedimentaryrocks, the extent to which pigmentation overprints primary paleomag-netic signals needs to be assessed on a case-by-case basis. Abrajevitchet al. (2014) analysed samples from the ~2.5 Ga Pilbara Print Stone(Fig. 4e, f) from Western Australia that lacks a detrital component, andprovides a rare opportunity to understand paleomagnetic recording as-sociated with pigment formation. The pigment in these rocks formedhydrothermally (Rasmussen et al., 2007), so its formation mechanismis not pertinent to our aim of considering diagenetic processes mainly

Fig. 4. Iron (oxyhydr-)oxide precipitation in oxic diagenetic environments. (a–c)Oxidation fronts related to fluidflow in the porousfluvial Hawkesbury Sandstone (Triassic), Sydney Basin,Australia. Mineralisation associatedwith flowof reduced iron in solutionwith precipitation of Fe3+-bearingminerals at oxidation fronts that: (a) have travelled through the rock and tran-sect bedding planes (scale: Australian 20 cent coin is 28.5 mm across), (b) were restricted to fluid flow in fractures, withmineral precipitation around the termination of a fracture, wheregrey colours along the fracture indicate that iron remains in a reduced state (scale: Australian 50 cent coin is 31.5mmacross), and (c)flowed along both fracture andbedding planes (scale:Australian 50 cent coin). Typical iron (oxyhydr-)oxide minerals in the Hawkesbury Sandstone are limonite, goethite and haematite. Paleomagnetic studies indicate that suchmineralisation caused late Cenozoic to Pleistocene remagnetisations (Bishop et al., 1982; Pillans, 2003). The observed banding of iron (oxyhydr-)oxides is common in chemical systemsundergoing a single precipitation event (Liesegang, 1896). (d–f) Haematitemineralisation in the Hamersley Basin,Western Australia, from (d) the 3.46 GaMarble Bar Chert (Hoashi et al.,2009), and (e, f) ‘print stone’ from the 2.5 Ga Mount McRae Shale formation (Abrajevitch et al., 2014). The haematite in these examples formed via reaction of Fe2+-rich hydrothermalfluids andO2-rich seawater; areaswith reduced and oxidised iron are evident in all of the outcrops. The haematite in (e, f) occurs as newsprint and uniformpigment types (see Abrajevitchet al. (2014) for details). The outcrop in (d) is ~2 m across. The scale in (e) is an Australian 50 cent coin. In (f), the pigment loop with arrow on the right-hand side is ~40 mm across.


below ~50 °C. However, pigment precipitation in the Print Stone can belinked paleomagnetically to known episodes of regional fluidmigration,with precipitation occurring over long enough periods of time to recordnormal and reversed paleomagnetic polarities. The different pigmenttypes have contrasting magnetic properties, but some samples containboth pigment types. Overlappingmagnetic properties for these pigmenttypes makes it difficult to discriminate between the pigments on thebasis of magnetic properties, although the recorded paleomagnetic di-rectionsmake it straightforward to discriminate between them. Overall,Abrajevitch et al. (2014) demonstrated short-range variability in the oc-currence of different pigment types that indicates a strong local controlon pigment formation. This raises the possibility that a DRM can survivein red beds even in the presence of CRM-carrying pigments in localpockets, as suggested by Liu et al. (2011). In other cases, red beds withmagnetic properties controlled by detrital haematite have been demon-strated to preserve robust paleomagnetic records of short-term geo-magnetic secular variation (Kruiver et al., 2000). In yet other cases,earlymagnetisations have been demonstrated to be carried by both pig-mentary and detrital haematite in the same samples (Iosifidi et al.,2010). These contrasting results illustrate the need for careful discrimi-nation between different haematite components in paleomagneticstudies.

4.2.2. Nitrogenous magnetic mineral diagenesisThe nitrogenous diagenetic zone is often not explicitly discussed as

being important formagneticmineral diagenesis. However, upward dif-fusing Fe2+ is oxidised at the base of the nitrogenous zone, which rep-resents the Fe3+/Fe2+ boundary (Froelich et al., 1979). Also,magnetite-producing magnetotactic bacteria are usually nitrate re-ducers that live around the nitrogenous-ferruginous transition zone(Bazylinski and Moskowitz, 1997; Schüler and Baeuerlein, 1998; Flieset al., 2005) where both dissolved Fe2+ and an oxidant (i.e., nitrate)are present and counter-diffuse (Fig. 5). The nitrogenous zone is, there-fore, a key zone for iron oxide formation.

Recent work suggests that the nitrogenous zone might be addition-ally important for Fe redox cycling and that magnetic mineral diagene-sis could be more complex than was previously assumed. Magnetitebiomineralisation has so far been attributed to two main pathways, asdescribed below, which give rise to formation of intracellularmagnetiteproduced by magnetotactic bacteria and extracellular magnetite pro-duced by dissimilatory iron-reducing bacteria. Miot et al. (2014) pro-posed a further route for magnetite biomineralisation by anaerobicnitrate-reducing Fe2+-oxidising bacteria. The bacteria promote trans-formation of green rust (Fe2+-Fe3+ hydroxides) to lepidocrocite to ex-tracellular magnetite. The magnetite has stable single domain (SD)

Fig. 5. Schematic representation of stratified chemical environments inwhichmagnetotactic bacteria live. Left: Concentration of dissolved species in thewater column or pore fluidswithnitrogenous conditions giving way with depth to ferruginous and sulphidic conditions. The so-called oxic–anoxic transition zone (Kopp and Kirschvink, 2008; vertical line with arrows)corresponds to the ferruginous zone of Canfield and Thamdrup (2009). Centre: Chemical environments in whichmagnetite- and greigite-producingmagnetotactic bacteria live, with im-ages of magnetite-producing bacteria (above; from Lin et al., 2013) and greigite-producing bacteria (below; from Lefèvre et al., 2011). Right: Magnetite stability in the respective parts ofthe chemical zonation. Modified from Kopp and Kirschvink (2008).


magnetic properties that are ideal for paleomagnetic recording. Themagnetite-producing reactions documented by Miot et al. (2014)were obtained in laboratory culture, and it has yet to be demonstratedthat they occur in natural environments. Muchwork, therefore, remainsto assess the possibility of magnetite biomineralisation in the nitroge-nous diagenetic zone.

4.2.3. Ferruginous magnetic mineral diagenesisFe3+ is the second most abundant electron acceptor, after sulphate,

in marine sediments (Reeburgh, 1983), and is often the most abundantelectron acceptor in freshwater environments (Nealson and Myers,1992; Nealson and Saffarini, 1994). Iron, therefore, plays an extremelyimportant role in organic matter degradation, with the principal out-comebeing reductive dissolution of iron (oxyhydr-)oxides.Many bacte-rial species reduce iron and can do so in environments that range fromoxic to strictly anoxic. Iron-reducing bacteria are important for themag-netic mineralogy of sediments because their metabolism gives rise totwo major effects: first, they cause dissolution of reactive iron-bearingminerals (e.g., Kostka and Nealson, 1995); second, once dissolved ironbecomes available in pore waters, the type of mineral that then precip-itates (e.g., ferrimagnetic versus paramagnetic) controls the preservedmagnetisation. As outlined above, the minerals that are most reactiveto dissolution are hydrous ferric oxide, ferrihydrite and lepidicrocite(Poulton et al., 2004). Theseminerals are negligibly or weakly magneticand, therefore, their loss through dissolution does not have a major ef-fect on the magnetisation of sediments. As a result, they have largelybeen ignored in paleomagnetic studies. Their environmental impor-tance in iron cycling has been increasingly recognised in recent years,which has led to an upsurge in magnetic investigations of theseminerals particularly with respect to formation and transformationpathways of magnetic minerals to better understand environmentalprocesses (e.g., Michel et al., 2010; Liu et al., 2012a; Till et al., 2014).

Given that dissolution of hydrous ferric oxide, ferrihydrite andlepidicrocite does not significantly affect themagnetic properties of sed-iments, the most important process for the sedimentary magnetic re-cord in the ferruginous zone is magnetite biomineralisation. Two setsof processes give rise to magnetite biomineralisation in this zone andhave been referred to as biologically inducedmineralisation and bound-ary controlled mineralisation (Lowenstam, 1981; Lowenstam andWeiner, 1989; Moskowitz, 1995). Biologically induced mineralisationrefers to mineral formation where an organismmodifies its local extra-cellular microenvironment to create conditions suitable for inorganicchemical precipitation of minerals by reaction of extraneous ions withmetabolic products extruded across a cell wall. In contrast, boundarycontrolled mineralisation occurs when an organism exerts intracellularcontrol on mineral precipitation, which produces functional materialswith specific crystal chemical properties. Magnetotactic bacteria arethe most important microorganisms that produce magnetite throughboundary controlled mineralisation (e.g., Bazylinski and Frankel, 2004;Faivre and Schüler, 2008; Kopp and Kirschvink, 2008), while dissimila-tory iron-reducingbacteria are themost important example ofmicrobesthat produce magnetite through biologically induced mineralisation(e.g., Lovley, 1991; Lovley et al., 2004). In this context, dissimilation re-fers to the use of iron as a terminal electron acceptor for anaerobic res-piration for purposes other than intracellular assimilation. In thefollowing treatment, these contrasting types of biogenic magnetite arediscussed, respectively, in relation to their short- and long-term effectson the sedimentary magnetic record.

4.2.3.1. Magnetite magnetofossils in the sedimentary record. Intracellularbiomineralisation of magnetite by magnetotactic bacteria results inwell-controlled magnetic mineral stoichiometry, grain size, grainmorphology, and magnetosome chain structure (e.g., Bazylinski andFrankel, 2004; Faivre and Schüler, 2008; Kopp and Kirschvink, 2008).Magnetotactic bacteria in modern environments are mainly sampledfrom the water column or sediment–water interface, in which case

they are not a product of diagenesis. However, somemagnetotactic bac-terial populations have been reported to live within the surface mixedlayer of sediments (e.g., Petermann and Bleil, 1993; Flies et al., 2005;Pan et al., 2005; Jogler et al., 2010; Mao et al., 2014a, 2014b) or deeper(e.g., Tarduno andWilkison, 1996; Tarduno et al., 1998), which, strictlyspeaking, means that the biomineralised magnetosomes formed duringearliest burial (Fig. 6) so that they should be considered in discussions ofmagnetic mineral diagenesis. Manymagnetotactic bacteria derive ener-gy for metabolism by living within sharp redox gradients near the so-called oxic–anoxic interface or oxic–anoxic transition zone (Bazylinskiand Frankel, 2004; Faivre and Schüler, 2008; Kopp and Kirschvink,2008). In the scheme shown in Fig. 1, this zone is equivalent to the ni-trogenous–ferruginous transition (Fig. 5). Magnetotactic bacteria are,therefore, often intimately involved in diagenetic iron and carbon cy-cling to an extent that they should be considered in discussions of sed-imentary diagenesis.

Magnetotactic bacteria have worldwide distribution and are consid-ered to be ubiquitous in aquatic environments (e.g., Blakemore, 1982;Lefèvre and Bazylinski, 2013). After their discovery (Bellini, 1963a,1963b; Blakemore, 1975), they were, therefore, widely expected to bean important source of SDmagnetite that carries stable sedimentary pa-leomagnetic records (e.g., Kirschvink, 1982; Petersen et al., 1986; Stolzet al., 1986, 1990; Chang et al., 1987; Vali et al., 1987). However, asdiscussed below in relation to sulphidic diagenesis, a key issueconcerning the long-term preservation potential of biogenic magnetitewas not well appreciated: reductive dissolution of magnetite is ubiqui-tous in sulphidic environments below the ferruginous-sulphidic transi-tion (Karlin and Levi, 1983, 1985; Canfield and Berner, 1987; Karlin,1990a, 1990b; Leslie et al., 1990a, 1990b; Channell and Hawthorne,1990). Thus, when magnetite magnetofossils are buried and eventuallysubjected to sulphidic diagenesis, biogenic magnetite will dissolve, andiron sulphides will form (Berner, 1984).

Throughoutmy career, I have consideredmagnetotactic bacteria andthe magnetofossils that they produce to be interesting but largely irrel-evant in terms of the sedimentary paleomagnetic record. This view issupported by the fact that, until recently, only a small handful of reliablemagnetite magnetofossil identifications had been documented inthe pre-Quaternary geological record (Kopp and Kirschvink, 2008).However, recent application of techniques that are suited to identifyingintact SD magnetosome chains produced by magnetotactic bacteria(e.g., Weiss et al., 2004; Kopp et al., 2006; Egli et al., 2010), in additionto transmission electron microscope (TEM) observations, has led tore-evaluation of the paleomagnetic importance of magnetitemagnetofossils in the geological record (e.g., Kopp et al., 2007, 2009;Yamazaki, 2008, 2009, 2012; Abrajevitch and Kodama, 2009, 2011;Roberts et al., 2011a, 2012, 2013a; Chang et al., 2012; Larrasoaña et al.,2012; Yamazaki and Ikehara, 2012). Widespread preservation ofmagnetite magnetofossils throughout thick pelagic marine carbonatesequences has been attributed to the presence of thick nitrogenous–fer-ruginous zones where low organic carbon inputs gave rise to diageneticconditions that were never sulphidic and hence corrosive to biogenicmagnetite (Roberts et al., 2011a; Yamazaki and Shimono, 2013). Thus,whilemanymagnetotactic bacteria species livewithinmicroaerobic en-vironments (Blakemore et al., 1985) or near the nitrogenous–ferrugi-nous transition (e.g., Bazylinski and Frankel, 2004; Faivre and Schüler,2008; Kopp and Kirschvink, 2008) (Fig. 5), their preservation and con-tribution to long-term paleomagnetic recording indicates that theyhave never undergone reductive diagenesis (Roberts et al., 2011a;Yamazaki and Shimono, 2013). Evaluation of a large catalogue of sedi-ment samples resulted in recognition of widespread magnetitemagnetofossil preservation in ancient lake sediments and variable ma-rine environments, including mixed terrigenous-carbonate environ-ments, pelagic carbonates, and clay-rich pelagic sediments (Robertset al., 2012). It is expected that magnetite magnetofossils will be identi-fied more routinely in the geological record and will prove to be awidespread contributor to paleomagnetic signals. However, burial of


sediments into the sulphidic diagenetic zone will cause magnetite dis-solution (Karlin and Levi, 1983; Canfield and Berner, 1987; Karlin,1990a, 1990b), which will limit geological preservation of magnetitemagnetofossils.

For sediments that contain magnetofossils, the key questionconcerning the recorded paleomagnetic signal relates to the depth atwhich the magnetotactic bacteria lived (Fig. 6). If they lived within thewater column or within the surface mixed layer of the sediment, theirpost-mortem remains should align passively with the ambient geomag-netic field (Ouyang et al., 2014), much like detrital particles that recorda post-depositional remanent magnetisation (PDRM). However, if theylived below the surface mixed layer as suggested by Tarduno andWilkison (1996) and Tarduno et al. (1998), magnetofossils will contrib-ute to a biogeochemical remanentmagnetisation (Roberts et al., 2013a;Larrasoaña et al., 2014). Further work is needed to test these possibili-ties. Initial efforts suggest that the alignment efficiency of livingmagnetotactic bacteria in surface sediments is low (b1%), which is con-sistent with the typically poor magnetisation efficiency of sediments(Mao et al., 2014a, 2014b). Magnetite is not the only magnetic mineralproduced by magnetotactic bacteria; greigite can also be produced(e.g., Farina et al., 1990; Mann et al., 1990; Bazylinski et al., 1993;DeLong et al., 1993; Pósfai et al., 1998a; Reitner et al., 2005; Kasamaet al., 2006; Lefèvre et al., 2011; Chen et al., 2014). The possibility of bio-genic greigite preservation in the geological record is discussed below inSection 4.2.4.4.

4.2.3.2. Extracellular magnetite in the sedimentary record. Authigenicmagnetite is produced as an extracellular by-product of metabolism ofdissimilatory iron-reducing bacteria (Lovley et al., 1987). For example,Geobacter metallireducens (formerly known as GS-15) anaerobicallycouples organic matter oxidation to reduction of Fe3+, which inducesextracellular precipitation offine-grainedmagnetite (Lovley, 1991). Un-likemagnetotactic bacteria, lack of strict biological control on the extra-cellular crystallisation process results in particles that lack uniquemorphology and that have broad grain size distributions. Laboratory-cultured G. metallireducens has been documented to produce 5000times as much magnetite by weight as an equivalent biomass ofmagnetotactic bacteria in the laboratory (Lovley, 1991). Dissimilatory

iron-reducing bacteria could, therefore, be an important source offine-grained magnetite, and have been suggested to be important forthemagnetisation of sediments (Stolz et al., 1990). It is thought that dis-similatory iron reducers catalyse nearly all iron reduction in ferruginousenvironments (Lovley, 1993). Extracellular magnetite production is,therefore, likely to occur in any ferruginous setting, including deep-sea environments. For example, Wu et al. (2011) documented magne-tite production from Shewanella piezotolerans WP3, which was isolatedfrom sediments at water depths of ~1900m, althoughWu et al. (2013)suggested that extracellularmagnetite production is inhibited at hydro-static pressures N50 MPa (equivalent to water depths of 5000 m).

Despite the widespread occurrence of dissimilatory iron reducers,most particles produced by G. metallireducens occur in the magneticallyunstable ultrafine superparamagnetic (SP) state at room temperature(Moskowitz et al., 1993), so that only a small proportion of such parti-cles is likely to be large enough to be paleomagnetically important.The role of this type of magnetite in the sedimentary paleomagnetic re-cord is difficult to assess because of the broad grain size distributionsproduced and the lack of specific morphologies. Additionally, the pres-ervation potential of extracellular magnetite has been suggested to below (e.g., Li et al., 2009) because the large surface area of suchfinegrainsmakes them a likely substrate for wide-ranging physical, chemical, andbiological reactions (e.g., Navrotsky, 2000). Li et al. (2009) documentedoxidative degradation of ~13 nm extracellular magnetite under anoxicstorage over a 5-year period, which they suggested as an explanationfor the rare geological preservation of extracellular magnetite. Theyalso argued that anaerobic oxidation ofmagnetite contributes to furtherdiagenetic iron cycling. This suggestion is supported by the paucity ofdocumented cases of extracellular magnetite preserved in sediments.The best demonstrated cases involve modern rather than ancient sedi-ments (Gibbs-Eggar et al., 1999; Maloof et al., 2007). Synchrotron anal-yses indicate that extracellular magnetite has similar Fe cation siteoccupancies to abiogenic magnetite, but with oxygen deficiency dueto formation of the magnetite in oxygen-poor environments aroundthe nitrogenous–ferruginous transition (Coker et al., 2007). Other stud-ies indicate that both extracellularly producedmagnetite and intracellu-lar magnetite magnetosomes contain a higher Fe2+ concentration thanabiogenicmagnetite (Carvallo et al., 2008; Lamet al., 2010).While these

Fig. 6. Occurrences of magnetotactic bacteria living within sediment in a laboratory microcosm. (a) A 1 L microcosm with (b) measured profiles of O2 concentration (blue crosses) andmagnetotactic bacterial populations (cells/μl) within the surface sediment after the sediment had stabilised for 7 days. The bacteria include Magnetotacticum bavaricum (green squareswith error bars) and cocci (orange circles with error bars). Modified from Mao et al. (2014a).


mineralogical variations likely reflect environmental conditions duringmineralisation, it is unknownwhether the oxygen deficiency or relativeenrichment in Fe2+ with respect to Fe3+ confers any greater resistenceto dissolution once the particles are buried within sulphidic environ-ments. In a freshwater setting, Kodama et al. (2013) reported thatsome magnetite magnetofossils persisted to greater depths after com-plete dissolution of most magnetosomemorphologies had occurred, al-though the concentration of relict magnetofossils was small anddiagenetic dissolution became more pervasive with depth. Overall, itis unlikely that extracellular or intracellular magnetites will survivesulphidic diagenesis even if they are oxygen deficient and relativelyenriched in Fe2+ with respect to Fe3+.

In contrast to the suggestion that extracellular magnetite pro-duced by dissimilatory iron-reducing bacteria is poorly preservedin the geological record, several lines of evidence point to the preser-vation of SP particles even after burial in sulphidic diagenetic envi-ronments. Tarduno (1995) documented significant concentrationsof SP particles in diagenetically reduced pelagic carbonates. Extracel-lular magnetite could explain the presence of SP particles in thesesediments. Further possibilities are discussed below in relation tosulphidic diagenetic environments. The difficulty in understandingthe origin of such ultrafine particles is that multiple processes cangive rise to particles in this size range. Egli (2004a, 2004b, 2004c)recognised this difficulty and, therefore, grouped particles producedby dissimilatory iron-reducers with detrital and pedogenic magne-tite because of their similar magnetic coercivity distributions. Amag-netically non-interacting low-coercivity stable SD component isbeing identified increasingly in modern and ancient sediments; ithas been attributed to extracellular magnetite, although definitiveevidence is lacking (Egli, 2004a, 2004b, 2004c; Ludwig et al., 2013;Heslop et al., 2014). Much more work is needed to understand thecontribution of extracellular magnetite to the magnetic record ofsediments. The importance of magnetite magnetofossils for themagnetisation of sediments was not adequately appreciated untilappropriate tools were developed to recognise their presence(Roberts et al., 2012). It is possible that the same is true of extracel-lular magnetite. Thus, instead of dismissing the possible importanceof extracellular magnetite because of lack of widespread documenta-tion of its presence, renewed effort is needed to understand the ex-tent of its preservation in the geological record.

4.2.4. Sulphidic magnetic mineral diagenesisIn sulphidic sediments, dissolution of magnetite and haematite,

which are not among the most reactive iron-bearing minerals(Poulton et al., 2004; Table 3), becomes ubiquitous (Canfield andBerner, 1987). Dissolved Fe2+ released from detrital iron-bearing min-erals then reacts with dissolved H2S, which is a by-product of sulphatereduction, to form sedimentary iron sulphides, particularly pyrite(Berner, 1984). Dissolution of detrital magnetite and haematite insulphidic diagenetic environments, and replacement by paramagneticpyrite, which does not carry a permanent magnetisation, can destroythe primary paleomagnetic record (Karlin and Levi, 1983, 1985;Channell and Hawthorne, 1990; Karlin, 1990a, 1990b; Leslie et al.,1990a, 1990b; Roberts and Turner, 1993; Robinson et al., 2000; Rowanet al., 2009). This process is ubiquitous in continental margin marinesediments with high organic carbon contents. If the rate of Fe2+ supplyexceeds that ofH2S production, intermediate iron sulphides that form asprecursors to pyrite, such as mackinawite and ferrimagnetic greigite(Berner, 1984; Roberts and Turner, 1993; Fu et al., 2008), can be pre-served (Kao et al., 2004). Greigite can form early (e.g., Cutter andKluckhohn, 1999; Reynolds et al., 1999; Vasiliev et al., 2008; Blanchetet al., 2009; Hüsing et al., 2009; Nowaczyk et al., 2012) or it can growlater and remagnetise the host sediment (Roberts and Weaver, 2005)depending on the timing of availability of the necessary reactants.Distinguishing between these possibilities is fundamental to paleomag-netic analysis of sediments that have undergone sulphidic diagenesis. In

general, sulphate-reducing diagenetic environments are destructive topaleomagnetic recording, and oxic to ferruginous diagenetic conditionsaremore likely to preserve primary paleomagnetic signals. Severalmag-netically important processes that occur in the sulphidic zone arediscussed below.

4.2.4.1. Iron sulphide formation pathways. Iron sulphide formation path-ways are discussed in detail in several review papers (Rickard et al.,1995; Schoonen, 2004; Rickard and Luther, 2007). When consideringsedimentary organic matter decomposition and authigenic mineralformation under sulphidic conditions, the major reactions of interestare the pathways of hydrogen sulphide formation via organoclasticsulphate reduction:

2CH2Oþ SO2−4 ¼ H2Sþ 2HCO−

3 ; ð2Þ

or via anaerobic oxidation of methane (AOM):

CH4 þ SO2−4 ¼ HCO−

3 þHS‐ þ H2O: ð3Þ

The rate of these reactions is controlledmainly by the concentration andreactivity of decomposable organic matter (Berner, 1981). Dissolvedsulphide produced by sulphate reduction will then react with dissolvedFe2+ (produced by dissolution of detrital iron-bearingminerals via ironreduction) or directly with iron-bearing minerals to form iron sul-phides. Sulphidisation of Fe2+ in solution, via reaction with H2S, causesformation of sedimentary iron sulphides, starting with ironmonosulphides (Berner, 1970):

Fe2þ þH2S aq:ð Þ ¼ FeS sð Þ þ 2Hþ: ð4Þ

Several pathways have been suggested for sulphidisation reactions thatultimately produce pyrite from iron monosulphide and greigite precur-sors via reaction with sulphur species (H2S, HS−, Sn2−) that enable sul-phur addition (e.g., Schoonen and Barnes, 1991):

FeS sð Þ þH2S aqð Þ ¼ FeS2 sð Þ þH2 gð Þ; ð5Þ

FeS sð Þ þHS– þ H2O lð Þ ¼ FeS2 sð Þ þH2 gð Þ þ OH–; and ð6Þ

FeS sð Þ þ S2−n ¼ FeS2 sð Þ þ S2−n−1: ð7Þ

Other pathways involve FeS conversion to greigite or pyrite via iron loss(e.g., Lennie et al., 1997; Pósfai et al., 1998b). Debate continues aboutthese pathways (Schoonen, 2004). Regardless, the most important im-pact of sulphidisation on magnetic minerals is the reductive dissolutionof iron oxides (Canfield and Berner, 1987), which leads to formation ofparamagnetic pyrite and to progressive removal of detrital magneticminerals that partially or completely destroys the primary paleomag-netic signal. If pyritisation is arrested, authigenic growth of the ferri-magnetic intermediate iron sulphide, greigite, can give rise to a stablesecondary paleomagnetic signal. The key issue to resolve with regardto the fidelity of the paleomagnetic signal carried by greigite concernsthe timing of its formation, which can be highly variable depending onthe availability of dissolved iron and sulphide (Roberts and Weaver,2005).

Equilibrium thermodynamic considerations led to the conclusionthat greigite will be rare in the geological record (Berner, 1970, 1984)because pyrite rather than greigite is the stable iron sulphide overmuch of the range of Eh–pH values expected in sulphidic sediments(Berner, 1970, 1984). Widespread recognition of greigite, largelythrough its magnetic properties (e.g., Snowball and Thompson, 1988;Snowball, 1991; Horng et al., 1992; Roberts and Turner, 1993;Reynolds et al., 1994; Roberts, 1995; Fu et al., 2008; Roberts et al.,2011b), has led to greigite being reported routinely in sulphidic sedi-ments. Failure of equilibrium thermodynamics to explain paleomagnet-icallymeaningful greigite concentrations in the geological record results


from under-appreciation of the importance of kinetic considerations,where the rate of Fe2+ supply and H2S production appear to be themost important determinants for ferrimagnetic greigite preservation(Kao et al., 2004; Schoonen, 2004). These factors are controlled by theavailability of reactive iron and metabolisable organic matter.

Most laboratory studies of iron sulphide formation have excludedthe bacterial consortia that live in sulphidic sedimentary environments.It is now recognised that bacterial sulphate reduction plays an impor-tant role in iron sulphide formation beyond the traditionally assumedrole of providing H2S for sulphidisation reactions (Schoonen, 2004).For example, Donald and Southam (1999) demonstrated that sulphur-bearing amino acids from the cell walls of sulphate-reducing bacteriaare incorporated into iron sulphides (which provides support for a sul-phur addition pathway for iron sulphide formation). They also proposedthat bacterial cell walls provide a nucleation site for iron sulphides,where their anionic surfaces are capable of binding Fe2+ (cf.Konhauser, 1998), for which Watson et al. (2000) provided photomi-crographic evidence (Fig. 7). Iron sulphides often also occur asmineralised replacements of organic matter (Fig. 8), including the inte-rior of microfossil chambers (Fig. 9) (e.g., Raiswell, 1982; Roberts andTurner, 1993; Stumm and Morgan, 1996; Roberts et al., 2005; vanDongen et al., 2007; Nowaczyk, 2011). Grimes et al. (2001) demonstrat-ed experimentally that iron sulphide precipitation starts on the cellwalls of plant matter. Microbial processes and availability of cellularsubstrates from a range of biological materials are clearly importantfor iron sulphide formation.

4.2.4.2. Greigite formation mechanisms. Greigite forms from an ironmonosulphide precursor (e.g., Schoonen and Barnes, 1991; Wang andMorse, 1995, 1996; Pósfai et al., 1998b; Benning et al., 2000;Schoonen, 2004),which has been identified as amorphous FeS or poorlyorderedmackinawite (Lennie et al., 1995). As indicated in reactions (5–7), several reaction pathways from a FeS precursor to pyrite continue to

be debated (Schoonen, 2004). Lennie et al. (1997) and Pósfai et al.(1998b) provided strong observational evidence for transformation ofmackinawite to greigite via iron migration within the crystal latticewith overall iron loss, although other pathways cannot be ruled out.While it is difficult to conclude in favour of any particular greigite for-mation pathway, several mineral formation mechanisms have beenproposed (Roberts andWeaver, 2005). All of the proposedmechanismsrequire sources of dissolved sulphide and iron,which can be available atany time during diagenesis through a range of processes (Roberts andWeaver, 2005). The resulting greigite can, therefore, have a range ofages, depending on the timing of its formation. For example, greigitehas been documented to form in the water column of sulphidic waterbodies, such as in anoxic fjords or within the Black Sea (Cutter andKluckhohn, 1999). In such settings, greigite would give rise to a paleo-magnetic signal that is equivalent to that acquired by detrital particles.Greigite has also been observed to form rapidlywithin various sedimenttypes (e.g., Pye, 1981; Pye et al., 1990; Reynolds et al., 1999; Blanchetet al., 2009; Nowaczyk et al., 2012), so that any paleomagnetic signalwould be equivalent to that carried by detrital particles. Studies of an-cient marine sediments have also provided evidence that greigite pro-vides a paleomagnetic signal that was acquired sufficiently early tocarry a reliable record of geomagnetic field variations (Vasiliev et al.,2007; Hüsing et al., 2009; Chang et al., 2014). In other cases, claims ofgeomagnetically reliable early remanence acquisition by greigite (Tricet al., 1991) are difficult to reconcile with petrographic observations ofmultiple generations of greigite formation (Roberts et al., 2005).

Analysis of modern sedimentary environments provides a powerfulmeans of testing the timing of greigite formation in sediments. Numer-ous authors have documented active greigite formation several metresto tens of metres below the sediment–water interface in continentalmargin marine sediments, with greigite formation occurring thousandsto hundreds of thousands of years after deposition (Kasten et al., 1998;Jørgensen et al., 2004; Liu et al., 2004; Neretin et al., 2004; Riedingeret al., 2005; Larrasoaña et al., 2007; Fu et al., 2008; Rowan et al., 2009;Riedinger et al., 2014). This greigite formationmay still be early enoughto make the measured paleomagnetic signal useful for tectonic recon-structions, but it will complicate magnetostratigraphic studies, andcompromise analysis of short-period geomagnetic fluctuations. Multi-ple studies provide evidence of remagnetisations that resulted fromlater diagenetic greigite formation (e.g., Thompson and Cameron,1995; Florindo and Sagnotti, 1995; Jiang et al., 2001; Roberts et al.,2005; Sagnotti et al., 2005a, 2010; Rowan and Roberts, 2005, 2006,2008; Larrasoaña et al., 2007; Florindo et al., 2007; Porreca et al.,2009; Lucifora et al., 2012; Aben et al., 2014). The timescales overwhich greigite forms make it important to have robust methods to as-sess the timing of greigite formation.

Roberts and Weaver (2005) documented five secondary mineraltransformations that can give rise to remagnetisations involvinggreigite. Thesemechanisms include: (1) neoformation of greigite on py-rite surfaces (Jiang et al., 2001); (2) neoformation of greigite withincleavages of detrital sheet silicate grains (Jiang et al., 2001);(3) neoformation of greigite on surfaces of authigenic clays (Jianget al., 2001); (4) greigite growth on surfaces of siderite (Roberts andWeaver, 2005; Sagnotti et al., 2005a); and (5) greigite growth on sur-faces of gypsum (Roberts and Weaver, 2005; Florindo et al., 2007). Incases 1–4, the parent mineral provides an iron source in the presenceof dissolved sulphide for secondary sulphidisation reactions that pro-duce greigite. For example, greigite formation from siderite (case 4) isgiven by the following reaction (Krupp, 1991):

3FeCO3 þ 4H2S ¼ Fe3S4 þ 3H2CO3 þ H2: ð8Þ

In case (5), gypsum can form on the surfaces of pyrite nodules, whichrequires oxidation of earlier-formed pyrite due to non-steady statediagenesis in diagenetic environments that changed from sulphidicto oxic (e.g., in coastal plain sediments subjected to large-

Fig. 7. Transmission electron micrograph of the cell of a sulphate-reducing bacteriumgrown in a chemostat and encrusted by iron sulphide minerals (modified from Watsonet al. (2000)). The anionic surfaces of bacterial cell walls are proposed to be capable ofbinding Fe2+ (Konhauser, 1998) to provide a nucleation site for iron sulphides.


amplitude sea level changes; Florindo et al., 2007). When the sedi-ments later returned to reducing conditions, sulphate within gyp-sum provides a source of sulphate that can be reduced to sulphidethat can then react with dissolved iron to form greigite and pyrite(Roberts and Weaver, 2005).

In case (2), the half-life of sheet silicate reaction with sulphide is120,000 years (Canfield et al., 1992) (Table 3) when in contact with1 mM sulphide concentrations (i.e., typical of pore waters when mostsulphate has been consumed). In such cases, iron sulphide formationis limited by the availability of reactive iron; slow reaction of iron in

Fig. 8.Macroscopic iron sulphide nodules that have remineralised large-scale sedimentary organic matter fragments. (a) Outcrop view of iron sulphide nodules with partially oxidisedsurfaces, and (b) iron sulphide nodules (with a fossil gastropod) that have weathered out of the outcrop (both from mid-Pleistocene marine sediments from the Valle Ricca sectionnear Rome, Italy; Florindo and Sagnotti (1995) and van Dongen et al. (2007)). These nodules contain large greigite concentrations and are strongly magnetic. (c) Greigite nodules thatare attracted to a handmagnet (fromPlio-Pleistocenemarine sediments exposed in the Erhjen-chi section, southwestern Taiwan (Jiang et al., 2001)). (d–f) Images at progressively highermagnification of an iron sulphide nodule with evidence of plant matter remineralisation; vertical structures represent replacements of vascular bundles. Individual pyrite framboids areevident in (f), which have subsequently been completely overgrown by pyrite that has filled and replaced the remaining vascular structures. These images are from a sample from earlyPleistocenemarine sediments, Vrica, Italy (Roberts et al., 2010). (g, h) Elongated framboids and non-framboidal pyrite aggregates that probably represent remineralisation of plant cellularmaterial (from early Pleistocenemarine sediments, Crostolo River, Italy (Roberts et al., 2005)). (i–l) Images at progressively highermagnification of an iron sulphide nodulewith evidenceof plant matter remineralisation; vertical structures represent replacements of vascular bundles (from nodules of the type shown in (a, b) from Valle Ricca, Italy; Florindo and Sagnotti(1995) and van Dongen et al. (2007)). The platey minerals are non-magnetic hexagonal pyrrhotite. In (k), a void within the polished surface reveals a cluster of greigite (finest-grained)and pyrite (coarser-grained) framboids that have been overgrown by pyrrhotite (platey texture). These framboids are encased in a well-preserved organic matrix. Pyrrhotite formationoccurred during a later sulphidisation event associated with anaerobic oxidation of methane (van Dongen et al., 2007). The polished section in (l) contains evidence of early diageneticgreigite (finest-grained) and pyrite (coarser-grained) framboids that have been overgrown by pyrrhotite. Images in (d–l) are back-scattered electron images obtained with a SEM.


sheet silicates can result in iron sulphide formation, as observed byCanfield et al. (1992). If greigite is preserved within sheet silicate layersas a result of slow reaction of iron with dissolved sulphide, it will havemajor implications for the paleomagnetic record of such sediments. Py-rite, greigite and monoclinic pyrrhotite have been observed commonlywithin silicate sheets (e.g., Fig. 10) in remagnetised sediments (Jianget al., 2001; Roberts and Weaver, 2005; Rowan and Roberts, 2006;Roberts et al., 2010). In the examples shown in Fig. 10a–c, the sheet sil-icates remain only partially sulphidised, which suggests thatsulphidisation has not continued for many half-lives. In contrast,

monoclinic pyrrhotite dominates the sheet silicates in Fig. 10d–e,which suggests that the pyrrhotite grew over prolonged periods togive rise to a late remagnetisation (Roberts et al., 2010). Canfield et al.(1992) noted that sheet silicates could become less reactive, with longerreaction half-lives, during ongoing sulphidisation. Mossmann et al.(1991) reported that only 30–60% of sheet silicate iron from Peru Mar-gin sediments was consumed over several million years of exposure topore water sulphide; Raiswell and Canfield (1996) estimated this sili-cate iron to have reaction half-lives of 2 to 3 million years. These obser-vations indicate that greigite and pyrrhotite growth on the reactive

Fig. 9. Microscopic iron sulphide aggregates that have formed in microenvironments by remineralisation of organic matter within microfossil chambers. (a–c) Images at progressivelyhighermagnification of polyframboidal iron sulphide aggregateswithin the chambers of a foraminifer. In (b) and (c), the finest-grained particles are greigite and coarser-grained particlesare pyrite. Some of the pyrite framboids are overgrown by euhedral pyrite. (d) Coarse-grained euhedral pyrite crystals within a microfossil chamber, (e) non-framboidal pyrite aggregatewithin a nannofossil chamber, and (f) a mixture of framboidal and non-framboidal pyrite within the chambers of a foraminifer. (g–i) Coarse-grained euhedral pyrite crystal aggregateswithin microfossil chambers. (j, k) Framboidal and non-framboidal greigite and pyrite within foraminiferal chambers. (l) Complete euhedral pyrite overgrowth of a microfossil. Images(a–g) are from early Pleistocene marine sediments from Crostolo River, Italy (Roberts et al., 2005), and (h–l) are from early Pleistocene marine sediments from Vrica, Italy (Robertset al., 2010). All images are backscattered electron images obtained with a SEM.


surfaces of sheets within sheet silicates has the potential to be widelyresponsible for remagnetisations.

4.2.4.3. Iron sulphide textures and evolution of porewaters during sulphidicdiagenesis. Iron sulphide textures can provide important informationconcerning early diagenetic sulphidisation. These textures includeframboidal and polyframboidal pyrite and greigite aggregates(Figs. 8g, h, k, l, 9a–c, and 11a–e), along with euhedral pyrite or massesof euhedral pyrite that have overgrown detrital or previously formedauthigenic minerals or different types of organic matter (Figs. 9d, g–i, land 11g–i). Environmental interpretations can be made from the se-quence in which authigenic iron sulphides formed using such texturalevidence, as described below.

Framboids are spheroidal aggregates of greigite or pyrite nano- ormicro-crystals in which these crystals have equal size. The termframboid is derived from the French for raspberry (framboise) due tothe raspberry-like appearance (e.g., Fig. 8k) of these sulphide aggregates(Rust, 1935). Greigite and pyrite are often found in polyframboidal clus-ters, with these sulphides also forming less regular masses betweenframboids (e.g., Love and Amstutz, 1966; Love, 1971; Morse andCornwell, 1987; Wilkin and Barnes, 1997; Jiang et al., 2001; Robertsand Weaver, 2005; Roberts et al., 2005, 2010; Rowan and Roberts,2006; Szczepanik and Sawlowicz, 2010). These types of textures are ev-ident in Figs. 8, 9 and 11. The origin of framboids is still debated, al-though several conclusions can be made. Framboidal morphology doesnot require biological activity (e.g., Farrand, 1970; Sweeney andKaplan, 1973; Wilkin and Barnes, 1996) even though high framboidaliron sulphide concentrations tend to be associated with organic matter(Fig. 11f). It has been argued that formation of greigite framboids is anessential precursor to framboidal pyrite (Sweeney and Kaplan, 1973;Wilkin and Barnes, 1997), although Butler and Rickard (2000)argued that framboid formation does not require a greigite intermedi-ate. This invalidates the suggestion of Wilkin and Barnes (1997) thatmagnetically-induced aggregation of ferrimagnetic greigite particles is

partially responsible for development of framboidal morphologies.Greigite framboids have nowbeenwidely observed, and the constituentgreigite nanocrystals are always finer grained than nearby pyritemicro-crystals (e.g., Ariztegui and Dobson, 1996; Jiang et al., 2001; Roberts andWeaver, 2005; Roberts et al., 2005, 2010; Rowan and Roberts, 2006).This observation is consistent with the formation process for framboidalpyrite proposed byWilkin and Barnes (1997), who suggested four con-secutive processes for pyrite framboid formation. (1) Nucleation andgrowth of iron monosulphide nanocrystals. (2) Reaction of thesenanocrystals to formgreigite. (3) Framboid growth through aggregationof greigite nanocrystals with uniform particle sizes. (4) Replacement ofgreigite framboids by pyrite. Banfield et al. (2000) presented evidencefor nanoparticle coarsening by self-assembly to form larger crystals viarotation so that self-assembling nanoparticles adopt parallel structuralorientations and aggregate through formation of appropriate chemicalbonds. Their observations were for iron oxyhydroxide self-assembly,but similar processes could operate for iron sulphide nanocrystalaggregation.

Regardless of formation processes, an important aspect of pyriteframboids is that constituent crystals in any framboid have uniform size(Love and Amstutz, 1966; Wilkin et al., 1996). This observation hasbeen argued to indicate that the crystals nucleated simultaneously (asmonosulphides) and that the crystals grew at a constant rate beforethey aggregated to form framboids (Wilkin et al., 1996). Wilkin et al.(1996) analysed pyrite framboids from a range of settings and deter-mined that framboids in sediments below euxinic (i.e., sulphidic) watersare on average smaller and less variable in size (5.0 ± 1.7 μm; 1σ) thanthose from sediments underlying oxicwaters (7.7± 4.1 μm). Additional-ly, in sediments underlying euxinic waters, b4% of framboids have sizesN10 μm, while 10–50% of framboids from non-euxinic settings havesizesN10 μm. This observationhas beenused to assess bottomwater ven-tilation (e.g., Passier et al., 1997;Wilkin et al., 1997; Böttcher and Lepland,2000) and suggests that the generally large framboid diameters (N10 μm)in Fig. 11 are indicative of diagenetic growth within sulphidic sediments

Fig. 10. Iron sulphide formationwithin sheet silicate minerals. Phyllosilicates have long reaction half lives of ~120,000 years with respect to sulphide (Canfield et al., 1992; Table 3). Minorpyrite formation within phyllosilicates from early Pleistocene marine sediments from (a) Crostolo River (Roberts et al., 2005) and (b, c) Vrica, Italy (Roberts et al., 2010). (d, e) Extensivepyrrhotite formation in silicate sheets in early Pleistocene sediments from Vrica (Roberts et al., 2010). This extensive iron sulphide formation provides evidence for prolongedsulphidisation. (f) Pyrite formationwithin sheet silicates inmodern sediments at 1.03mbelow the sediment–water interface from the Ría di Vigo, Spain (Mohamed et al., 2011). All imagesare backscattered electron images obtained with a SEM.


below an oxic water column. The framboid size relationship fordetermining ancient water column conditions has been questioned(e.g., Roychoudhury et al., 2003); nevertheless, the fact that framboidsize distributions do not vary with burial depth indicates that framboidformation occurs during earliest diagenesis (Wilkin et al., 1996). Thismakes it possible to assess the paragenetic sequence of authigenicminer-al formation in sulphidic sediments from petrographic observations.

Raiswell (1982) suggested that framboidal and euhedral pyrite growin a paragenetic sequence associated with progressive pore water evo-lution. Framboidal pyrite is expected to grow first when reactive ironis abundant and sulphide can be created rapidly by sulphate-reducingbacteria in a pore water system with free diffusive supply of seawatersulphate. With ongoing burial, pore water Fe2+ is progressivelyexhausted, organic matter becomes more refractory, and the diffusivedistance of seawater sulphate increases so that pore waters have amore restricted connection to seawater. The sulphate reduction ratethen slows, Fe2+ must diffuse over longer distances to any remainingorganic matter, where dissolved sulphide concentrations are higher,so that pyrite formation is slowed. Euhedral pyrite is then more likelyto form directly without a greigite precursor with slower kinetics

responsible for the change in pyrite texture (Raiswell, 1982). Euhedralpyrite forms without taking framboidal morphology (e.g., Fig. 9g–i),and occurs as overgrowths of framboidal pyrite (e.g., Figs. 9l and 11g–i). This paragenetic sequence is widely observed in sulphidic environ-ments (e.g., Raiswell, 1982; Wilkin et al., 1996; Roberts and Weaver,2005). The presence of euhedral pyrite, therefore, reflects progressiveearly diagenetic pore water evolution. Alternatively, it can reflect alater secondary sulphidisation event in a diagenetic system with re-stricted availability of sulphide and iron. There is no need to assumethe latter option unless there is direct petrographic observation of over-growths of different generations of authigenic minerals or appropriategeochemical evidence (Roberts and Weaver, 2005). For example, thereis no such evidence in Fig. 9 or Fig. 11. In contrast, framboids are exten-sively but variably overgrown by hexagonal pyrrhotite in Fig. 8j–l. Or-ganic geochemical evidence indicates that this sample underwent alater sulphidisation event associated with anaerobic oxidation of meth-ane (van Dongen et al., 2007), which probably gave rise to pyrrhotiteformation. Petrographic and geochemical analyses, therefore, can pro-vide important constraints on the relative timing of authigenic mineralformation events.

Fig. 11. Back-scattered electron SEM images of iron sulphide nodules. All images are from early Pleistocene marine sediments from Crostolo River, Italy (Roberts et al., 2005). (a, b) Low-magnification images of polyframboidal aggregates of pyrite and greigite, with less regular aggregates filling the spaces between framboids. These aggregates have replaced sedimentaryorganic matter fragments. (c) Close-up image of area in (b) in which individual iron sulphide crystals are resolved. The finest-grained particles are greigite, while all coarser particles arepyrite. Both greigite and pyrite occur as framboids; some of the pyrite framboids are overgrownwith euhedral pyrite, which results from progressive sulphidisation and pore water evo-lution (Raiswell, 1982). (d) Low- and (e) high-magnification views of an iron sulphide aggregatewith both greigite and pyrite present. (f) Schematic illustration of iron sulphide formationas a replacement of organic matter that decays by sulphate reduction and produces dissolved H2S that diffuses into surrounding Fe2+-rich pore waters to form iron sulphides (modifiedfrom Raiswell et al. (1993)). (g–i) Polyframboidal iron sulphide aggregates that have beenmore pervasively overgrown by euhedral pyrite (see text and Raiswell (1982) for explanation).


4.2.4.4. Greigitemagnetofossils in the geological record.Magnetotactic bac-teria most commonly live within the nitrogenous–ferruginous transitionzone in aquatic and sedimentary environments, with different types ofbacteria living within specific, narrow regions of the redox zonation(Fig. 5). Magnetite-producing bacteria are usually microaerophiles ornitrate-reducers and are associated with the nitrogenous part of the ni-trogenous–ferruginous transition at the top of the ferruginous zone(Bazylinski and Moskowitz, 1997; Schüler and Baeuerlein, 1998; Flieset al., 2005), while greigite-producing magnetotactic bacteria occurwithin the ferruginous-sulphidic transition at the base of the ferruginouszone (Bazylinski et al., 1993; DeLong et al., 1993; Chen et al., 2014).Greigite-producingmagnetotactic bacteria have been reported relativelywidely from sulphidic aquatic environments (e.g., Farina et al., 1990;Mann et al., 1990; Pósfai et al., 1998a; Reitner et al., 2005; Kasamaet al., 2006; Lefèvre et al., 2011; Chen et al., 2014), but have not yetbeen widely reported from the geological record. This is surprising be-cause, unlike iron oxides, greigite is stable under reducing diageneticconditions. In their relatively recent review of magnetofossil occur-rences, Kopp and Kirschvink (2008) stated that: “greigite magnetofossilshave not yet been identified in Quaternary sediments”. Two reports ofgreigite magnetofossils (Pósfai et al., 2001; Vasiliev et al., 2008) lackedthe definitive evidence required by the magetofossil scoring scheme ofKopp and Kirschvink (2008), andwere, therefore, classified as “putative”occurrences. Biogenic greigite has been identified positively withinsapropelic Holocene sediments from the Baltic Sea (Reinholdsson et al.,2013). Some Pliocene Mediterranean sapropels (Roberts et al., 1999)have similar magnetic properties to the Holocene Baltic sapropels(Reinholdsson et al., 2013), which suggests that biogenic greigite mayalso be the dominant magnetic mineral in these Mediterraneansapropels. Chang et al. (2014) provided the only demonstration to dateof greigite magnetofossil preservation in ancient sulphidic sediments(in Late Miocene sediments from the eastern European ParatethysSea). Identification of greigite magnetofossil chains is difficult becausegreigite magnetosomes have less regular crystal morphologies (Fig. 5)and less well aligned chain structures than magnetite magnetosomes(e.g., Farina et al., 1990; Mann et al., 1990; Pósfai et al., 1998a; Kasamaet al., 2006; Lefèvre et al., 2011). While definitive reports of greigitemagnetofossils in ancient sediments are still rare, it is reasonable to ex-pect them to be reported with increasing frequency in the coming years.

4.2.4.5. Extracellular greigite. In addition to extracellular magnetite pro-duced by dissimilatory iron reducers, extracellular greigite has alsobeen documented as a biologically induced mineral by-product ofsulphate-reducing bacteria (Watson et al., 2000) (Fig. 7). As is the casefor the dissimilatory iron reducer G. metallireducens, the greigite pro-duced by the sulphate reducers dominantly occurs in the magneticallyunstable ultrafine SP state (Watson et al., 2000). The sulphate reducersstudied byWatson et al. (2000) were grown in a chemostat with nutri-ent conditions that are not typical of natural environments, so it remainsto be tested how significant this type of greigite could be in the geolog-ical record. Nevertheless, magnetically significant concentrations of SPparticles have been documented in sulphidic sediments (Tarduno,1995; Rowan and Roberts, 2006; Rowan et al., 2009). Given thatultrafine-grainedmagnetite is likely to dissolve rapidly in such environ-ments, the source of SP particles proposed by Tarduno (1995) in suchenvironments could be greigite (Rowan and Roberts, 2006; Rowanet al., 2009). The widespread occurrence of SP greigite in sedimentsthat have been subjected to sulphate reduction has been attributed toinorganic mineral authigenesis starting with production of ultrafinenanoparticles that did not mature to form larger particles (Rowan andRoberts, 2006; Rowan et al., 2009). It is possible that initial nucleationof these nanoparticles resulted from extracellular greigite productionby sulphate-reducing bacteria as observed by Watson et al. (2000) orit could have resulted from arrested nanoparticle growth associatedwith inorganic mineral precipitation (Rowan and Roberts, 2006).Much remains unknown about the importance of extracellular greigite

in the geological record; more work is needed to understand thispossibility.

4.2.4.6. Pyrrhotite formation. While greigite is the most commonauthigenic magnetic iron sulphide mineral in sulphidic diagenetic envi-ronments, it is important to consider the conditions under which pyr-rhotite forms. Pyrrhotite occurs in various forms, the most importantof which for sediment magnetism is monoclinic pyrrhotite (Fe7S8),which is ferrimagnetic at ambient temperatures. Hexagonal (Fe9S10)and other forms of pyrrhotite are antiferromagnetic at room tempera-ture (O'Reilly et al., 2000), so they are not directly important in paleo-magnetism. Authigenic pyrrhotite has been widely claimed to beresponsible for early diagenetic sedimentary paleomagnetic signals(e.g., Linssen, 1988; Quidelleur et al., 1992; Mary et al., 1993; Robertsand Turner, 1993; van Velzen et al., 1993; Horng et al., 1998; Sagnottiet al., 2001; Channell and Stoner, 2002; Vasiliev et al., 2004; Kao et al.,2004). Such claims are not well founded because kinetic data suggestthat pyrrhotite formation will be extremely slow below ~250 °C(Lennie et al., 1995). So why has pyrrhotite been inferred to occur as aprecursor to pyrite in sulphidic sediments? This interpretation appearsto be the result of two errors. First, it appears to be linked to the infer-ence by Sweeney and Kaplan (1973) of the presence of hexagonal pyr-rhotite in laboratory syntheses up to 85 °C. However, Sweeney andKaplan (1973) did not definitively identify pyrrhotite in their syntheses,and subsequent work has failed to identify hexagonal pyrrhotite below180 °C;monoclinic pyrrhotite has never been identified in these synthe-ses (Schoonen and Barnes, 1991). Thus, it appears unlikely that mono-clinic pyrrhotite can grow fast enough in sediments to be a diageneticproduct during earliest burial. Second, it appears that the unconfirmedformation of antiferromagnetic hexagonal pyrrhotite as a precursor topyrite (Sweeney and Kaplan, 1973) has been misinterpreted as indica-tive of ferrimagnetic monoclinic pyrrhotite formation during early dia-genesis. This appears to be a case of mistaken identity and subsequentliterature reinforcement (see Horng and Roberts, 2006). Interpretationsof early ferrimagnetic pyrrhotite formation in paleomagnetic studiesare, therefore, unlikely to be correct. This evidence supports the obser-vation that authigenic monoclinic pyrrhotite is unknown in modernsulphidic sediments (Rickard et al., 1995).

Horng and Roberts (2006) demonstrated that detrital monoclinicpyrrhotite can be eroded frommetamorphic country rocks, transported,and deposited in depositional sinks during rapid flood events (Fig. 12a–c). They suggested that detrital rather than authigenic pyrrhotite is like-ly to be responsible for apparently reliable sedimentary paleomagneticsignals that are consistent with those recorded by detrital magnetite(e.g., Horng et al., 1998). Monoclinic pyrrhotite is produced whenpyrite-bearing sediments are subjected to metamorphism in regionalmetamorphic belts (Carpenter, 1974; Hall, 1986; Rochette, 1987). Ero-sion of these pyrrhotite-bearing host rocks fromgloballywidespread re-gional metamorphic belts could, therefore, make detrital pyrrhotite arelatively common carrier of sedimentary paleomagnetic signals(Horng and Roberts, 2006; Horng et al., 2012). Notwithstanding thepossibility that sedimentary pyrrhotite can have a detrital origin, slowformation of monoclinic pyrrhotite can occur (Fig. 12d–k). Its presenceis a clear indicator that paleomagnetic signals will be complex and willnot date to the time of deposition (e.g., Dinarès-Turell and Dekkers,1999; Weaver et al., 2002; Larrasoaña et al., 2007; Roberts et al.,2010). The contrasting petrographic context andmorphology of detrital(Fig. 12b, c) compared to authigenic pyrrhotite (Fig. 12d–k), and thehigh electron backscatter of iron sulphides compared to siliciclastic sed-iment components, makes it relatively straightforward to distinguishthese modes of pyrrhotite occurrence. Diagenetic pyrrhotite formationis discussed further in Section 4.2.5.

4.2.5. Methanic magnetic mineral diagenesisIf reactive sedimentary organic matter remains after pore water sul-

phate has been consumed, methanogenesis is the final step by which


organic matter is degraded. In freshwater, where sulphate concentra-tions are lower than inmarine environments, sulphate is exhausted rap-idly and methanogenesis occurs more rapidly (Capone and Kiene,1988). Methane production occurs as a by-product of the metabolismof strictly anaerobic archaea, where the terminal electron acceptor iscarbon. The most important known process associated with methanicdiagenesis for magnetic minerals relates to AOM (reaction (3) above).Sulphate reduction via AOM gives rise to consumption of pore watermethane and sulphate to depletion, where this reaction zone is referredto as the sulphate-methane transition (SMT). Consortia ofmethanotrophic archaea and sulphate-reducing bacteria (Hinrichset al., 1999; Boetius et al., 2000; Thiel et al., 2001; Knittel and Boetius,

2009) oxidise methane at the SMT, which produces hydrogen sulphide(Goldhaber, 2003; Burdige, 2006) as given in reaction (3) above(Murray et al., 1978; Devol and Ahmed, 1981; Niewöhner et al., 1998;Kasten and Jørgensen, 2000; Jørgensen and Kasten, 2006). AOM is mag-netically important because it provides a secondary, relatively mobile,source of hydrogen sulphide that can cause both reductive dissolutionof detrital iron oxides and formation of secondary ferrimagnetic ironsulphides. For example, distinctminima inmagneticmineral concentra-tion have been reported at the SMT in association with H2S release,magnetite dissolution, and pyrite formation (Garming et al., 2005;Riedinger et al., 2005; März et al., 2008). In other cases, peaks in mag-netic mineral content due to greigite formation have been reported at

Fig. 12. Detrital and authigenic monoclinic pyrrhotite from sediments. Detrital pyrrhotite from (a) a sediment trap and (b) within a polished section from surface sediments from a boxcore in the Taiwan Strait, with (c) a lower-magnification viewof graded beds from the box core associatedwith distinct gravitational settling events (due to typhoons). Images in (a–c) arefromHorng and Roberts (2006). The pyrrhotite is sourced from the regional metamorphic belt in the Taiwan Central Range (Horng and Roberts, 2006; Horng et al., 2012). Authigenic pyr-rhotite frommarine sediments is shown in (d–k). (d) Close-up and (e) wider-field views of authigenic pyrrhotite plates that have overgrown a siderite cement (grey) in sediments fromHydrate Ridge, offshore of Oregon (Larrasoaña et al., 2007). (f) Extensive pyrrhotite formation and (g) pyrrhotite overgrowths on siderite cement (grey) in early Pleistocene sedimentsfrom Vrica, Italy (Roberts et al., 2010). Pyrite framboids are evident within voids in the surface of the iron sulphide nodule in (f). (h–k) Nodular masses of intergrown pyrrhotite platesin early Miocene marine sediments from Sakhalin, Russia (from Weaver et al. (2002)). Authigenic pyrrhotite formation in all cases was probably associated with anaerobic oxidation ofmethane (see Larrasoaña et al., 2007). All images are backscattered electron images obtained with a SEM.


the active sulphidisation front where dissolved sulphide has diffuseddown from the SMT and reacted with upward diffusing Fe2+ (Fig. 1)to form greigite (Kasten et al., 1998; Jørgensen et al., 2004; Neretinet al., 2004; Horng and Chen, 2006; Larrasoaña et al., 2007). H2S forma-tion via AOM, and related iron sulphidemineral formation, occurs at thetop and base of the sulphidic zone; it is described here in relation tomethanicmagneticmineral diagenesis because itwould not occurwith-out methanogenesis.

In addition to the above processes, Fe2+ concentrations can increasewithin the methanic zone (Fig. 1), which has been suggested to occurdue to a coupling of AOM to Fe and Mn reduction. This process is re-ferred to as Fe- or Mn-mediated AOM (Beal et al., 2009; Sivan et al.,2011; Segarra et al., 2013; Riedinger et al., 2014; Egger et al., 2015).The microbial/metabolic pathways that drive iron reduction inmethanic sediments are strongly debated, but this diagenetic liberationof Fe2+ into pore waters below the sulphidic zone provides scope forsecondary formation of non-ferrimagnetic minerals such as siderite(FeCO3), which is a useful indicator of past methanogenesis (Berner,1980). Some authors have described the liberation of Fe2+ into porewaters below the sulphidic zone as a “cryptic sulphur cycle”(e.g., Holmkvist et al., 2011). If the liberated Fe2+ diffuses upward andencounters H2S, late iron sulphide formation becomes possible. Theseprocesses, therefore, have important implications for magnetic mineraldiagenesis and iron cycling in the methanic zone.

While dissolved sulphide production occurs at the SMT through theconventional steady-state diagenetic progression (Fig. 1), it is the po-tential mobility of methane through fracture and fault networks in tec-tonically active settings, which can happen at any time rather thansimply after burial to a certain depth, that makes AOM an under-appreciated but important process for understandingmagnetic mineraldiagenesis. While an association between methane-rich sedimentsor methane hydrates and greigite formation is well established(e.g., Housen and Musgrave, 1996; Horng and Chen, 2006; Musgraveet al., 2006; Enkin et al., 2007), analysis of sediment cores fromHydrateRidge, offshore of Oregon, clearly demonstrated greigite andmonoclinicpyrrhotite formation associated with methane venting (Larrasoañaet al., 2007). The variable timing of this activity adds complexity to pa-leomagnetic interpretation. Identicalmagnetic properties to those iden-tified by Larrasoaña et al. (2007) have been reported in studies fromsimilar geological settings (e.g., Roberts et al., 2010; Kars and Kodama,2015), which raises the possibility of the existence of tell-tale magneticsignatures associated with the former presence of methane hydrates.Organic geochemical analysis can enable detection of biomarker signa-tures of AOM. van Dongen et al. (2007) analysed iron sulphide nodulesfrom the Valle Ricca section, Italy, where Florindo and Sagnotti (1995)identified late remagnetisations, and found biomarker and isotopicsignatures that associate the iron sulphides with AOM. These resultsprovide the geochemical data needed to link remagnetisationsto AOM. Rowan and Roberts (2006, 2008) reported widespreadremagnetisations due to later diagenetic growth of greigite in Neogenesediments fromNewZealand and attributed them to tectonically drivenmethane migration and methane hydrate dissociation. AOM associatedwith non-steady state diagenesis, due to tectonic forcing of fluidsthrough sediments, provides a mechanism that is likely to be increas-ingly reported as responsible for remagnetisations associatedwithmag-netic iron sulphide minerals.

4.3. Magnetic mineral diagenesis in different sedimentary environments

4.3.1. Near-shore marine environmentsReactive organic carbon is usually abundant in near-shore marine

environments, the degradation of which causes organic matter diagen-esis to proceed rapidly into sulphidic conditions at depths of a few cen-timeters to tens of centimeters below the sediment–water interface. Insome settings, such as fjords, enclosed marine basins, or inland seas,where there is limited ventilation or exchange with oxygenated waters,

the water column can become sulphidic (e.g., Cutter and Kluckhohn,1999). Canfield and Berner (1987) conducted their classic study ofmag-netite diagenesis in coastal marine sediments from Long Island Sound,and concluded that magnetite dissolution will be ubiquitous in sedi-ments that support active sulphate reduction and H2S formation. InLong Island Sound, magnetite has half-lives that range from 50 to1000 years with dissolution occuring at shallow depths (tens of centi-meters below the sediment–water interface). The extent to whichmag-netite undergoes dissolution depends on several factors, including thesurface area of a magnetite particle, the concentration of dissolved sul-phide, and the time over which magnetite is in contact with sulphidicpore fluids. These factors, in turn, depend on the mineralogy and pre-depositional history of the sediment source(s), sedimentation rate, sul-phate reduction rate, and reactivity of the iron-bearingminerals depos-ited. Preservation of magnetite, if it happens at all, is most likely to beassociated with rapid burial, combined with high concentrations of re-active iron, which prevents build-up of pore water sulphide and totalmagnetite dissolution. Magnetite dissolution has now been observedin sulphidic diagenetic environments throughout the world. Paleomag-netic records from such environments are almost always destroyed,particularly in near-shore environments, although magnetite preserva-tion is possible (see Section 4.3.2).

A particularly clear example of the relationship between organic car-bon supply, sulphate reduction, magnetite dissolution, and iron sul-phide formation is provided by an onshore–offshore transect in themagnetically well studied river valley systems of NW Spain (Emirogluet al., 2004; Rey et al., 2005; Mohamed et al., 2011) that were drownedby sea level rise at the last glacial termination. Sedimentmagnetic prop-erties define a clear three-part zonation in the Ría de Vigo estuary, Spain(Fig. 13;Mohamed et al., 2011). The uppermost oxic zone ismagnetical-ly dominated by detrital Fe–Ti oxides, which undergo rapid down-coredissolution in an intermediate zone with magnetite dissolving beforehaematite, as expected from Table 3 (Poulton et al., 2004). Pyrite isabundant, alongwith low greigite concentrations, in this zone. The low-ermost sulphidic zone is almost completely depleted of magnetic min-erals. The upper two zones become thicker and occur at greater depthas organicmatter contents decreasewith distance from inshore settingsto deeper-water, less biologically productive environments at the estu-ary mouth (Mohamed et al., 2011). In organic-rich sediments of the Ríade Arousa estuary, beneath rafts used for shellfish aquaculture (withtotal organic carbon contents of up to 14%), magnetite has a half-lifeof less than 5 years (Emiroglu et al., 2004). If magnetite is depleted tonear-zero levels after 10 half-lives, the estimates of Canfield andBerner (1987) and Emiroglu et al. (2004) constrain the likely survivalduration of detrital paleomagnetic records in organic-rich near-shoremarine environments.

While it is expected that detrital magnetic minerals will undergorapid reductive dissolution at shallow depths within sediments innear-shore marine environments, the extent of this loss and the depthat which it occurs within the sediment will vary from location to loca-tion depending on the reactivity of organic matter and dynamic factorssuch as current strength, sediment resuspension and associated reoxy-genation of the sediment and organic matter (e.g., Rey et al., 2005).Zheng et al. (2010, 2011) andGe et al. (2015) analysedmagneticminer-al diagenesis in the inner shelf of the East China Sea,which is dominatedby sediment and organic carbon supply from the Yangtze River. Burialefficiency of organic carbon in fully marine environments tends to in-crease with sediment accumulation rate (e.g., Hedges and Keil, 1995;Burdige, 2005). In contrast, river-dominated continental margins havesignificant fractions of terrestrial organic matter, which is less reactivethan marine organic carbon (Burdige, 2005). Sulphate reduction ratesare, therefore, likely to be lower in such sediments. Additionally, organiccarbon is more effectively oxidised in highly dynamic fluid muds inthese settings, which can be deposited, resuspended and thenredeposited multiple times before final burial (Blair and Aller, 2012).Zheng et al. (2010, 2011) and Ge et al. (2015), therefore, found that


the diagenetic zonation of sediment cores from the inner shelf of theEast China Sea occurs over a greater sediment depth interval comparedto themarine carbon dominated sediments of the Ría de Vigo estuary inSpain (Mohamed et al., 2011). Likewise, spatial variability in organiccarbon supply, its reactivity, and extent of sediment remobilisation,mean that down-core magnetic zonations in such environments willbe less consistent than in the onshore–offshore transect of Mohamedet al. (2011) in the Ría de Vigo estuary. Neveretheless, diagenetic zona-tions in such sediments can provide a means of assessing relative varia-tions in organic carbon decomposition. Despite extensive diageneticmodification of the magnetic mineral assemblage in near-shore marinesediments, Zheng et al. (2010) reported relict magnetic particle assem-blages that record environmental signals associatedwith weak summermonsoon events.

Tidal flats from the intertidal zone represent a near-shore marineenvironment that is under-studied in terms of mineral magnetism.Prolonged greigite growth within tidal flat sediments of the Yangtze es-tuary has been associated with the refractory nature of slowlydegrading terrestrial organic matter and ongoing sulphidisation (Chenet al., 2015). While such sediments are not widely targetted for detailedmineral magnetic investigations, their sensitivity to sea level changemakes them a worthy target for environmental analysis.

4.3.2. Hemi-pelagic marine environmentsHemi-pelagicmarine sediments occur on continental shelves, slopes

and rises. The classic early studies of magnetite diagenesis of Karlin andLevi (1983, 1985) involved hemi-pelagic sediments offshore of Oregon.Magnetite dissolution has subsequently been widely detected in hemi-pelagic sediments around the world (e.g., Karlin, 1990a, 1990b; Leslieet al., 1990a, 1990b; Roberts and Pillans, 1993; Roberts and Turner,1993; Richter et al., 1999; Yamazaki et al., 2003; Liu et al., 2004;Garming et al., 2005; Riedinger et al., 2005; Dillon and Bleil, 2006;Kawamura et al., 2007; Rowan et al., 2009; Bouilloux et al., 2013;Chang et al., 2015). These studies provide a common picture of surfacesediments, generally with all stages of organic matter diagenesisrecognised in porewater profiles (Fig. 1). Oxic surface sediments gener-ally contain trace abundances of detrital iron oxide minerals, includingferric oxyhydroxides. Surficially oxidised magnetite particles will oftenhave irregular curved cracks on their surfaces (Fig. 3) due to latticemis-matches between themagnetite core and themaghemite rim (Petersenand Vali, 1987). As these maghemitised particles pass through the

ferruginous diagenetic zone, the surface maghemite skin is dissolved,which leaves a smaller remnant magnetite particle (Torii, 1997;Smirnov and Tarduno, 2000; Yamazaki et al., 2003; Kawamura et al.,2007; Yamazaki and Solheid, 2011).With ongoing burial, reactive mag-neticminerals undergomore pervasive dissolution once sulphide is pro-duced in porewaters, with thefinest particles dissolvingfirst, and pyritebecoming increasingly abundant. The depth atwhich themagneticmin-eral content declines precipitously depends on the organic carbon con-tent and sedimentation rate, and can vary significantly (e.g., ~30 cm,Karlin and Levi, 1983; and ~3.6 m, Kawamura et al., 2007), as illustratedfor an onshore–offshore diagenetic transect (Fig. 14) from the southernSea of Okhotsk, western Pacific Ocean (Kawamura et al., 2007).

The usefulness of paleomagnetic records from such diageneticallyreduced hemi-pelagic sediments depends on the extent to which detri-talmagnetic minerals survive dissolution. Hemi-pelagic sediments withhigh reactive iron contents are more likely than coastal sediments tohave moderate organic carbon contents and to be deposited rapidlyenough to avoid complete dissolution of the detrital magnetic mineralassemblage. For example, Plio-Pleistocene marine sediments fromsouthwestern Taiwan, which have been extensively sulphidised, con-tain mixtures of detrital magnetite, detrital pyrrhotite, and authigenicgreigite (Horng et al., 1992, 1998). The pyrrhotite was originallythought to have formed diagenetically, but it has subsequently beendemonstrated to be detrital in origin (Horng and Roberts, 2006; Hornget al., 2012). The pyrrhotite andmagnetite carry a consistent paleomag-netic record (Horng et al., 1998), as would be expected for detrital min-erals, while the greigite formed at variable times during diagenesis andgave rise to a contradictory polarity record (Jiang et al., 2001). Survivalof magnetite is likely to have arisen from the factors proposed byCanfield and Berner (1987): sedimentation rates were exceptionallyhigh, of the order of 1.5–3 m/kyr (Horng et al., 1998), and the marineclaystones deposited offshore of Taiwan are particularly Fe-rich (Kaoet al., 2004), probably due to erosion of Fe3+-rich tropical soils. In set-tings where detrital magnetite has been completely dissolved, the pa-leomagnetic record is interpreted to be recorded by relict magneticmineral assemblages (Hounslow et al., 1995; see Section 4.7).

Early studies of magnetic mineral diagenesis in hemi-pelagic sedi-ments inferred the possible presence of intermediate iron sulphidesalong with pyrite (e.g., Karlin and Levi, 1983). Assessing the presenceof greigite was difficult because its magnetic properties were poorlyknown until more recently (Roberts, 1995; Roberts et al., 2011b). Liu

Fig. 13. Illustration of diagenetic modification of magnetic properties with depth in sediment cores from the Ría di Vigo estuary, northwestern Spain. χ is the low-field magnetic suscep-tibility (a measure of magnetic mineral concentration) and ARM is the anhysteretic remanent magnetisation (a measure of fine (submicron) ferrimagnetic mineral concentration). Var-iable depths of the redox-related diagenetic transitions are caused by higher organic carbon contents in sediments from the inner estuary (gravity core ZV-1) with respect to the middle(core CGVIR-36) and outer estuary (CGPL-00-6). Colour: brown= oxic, olive green = ferruginous, and grey = sulphidic diagenetic conditions. Modified from Mohamed et al. (2011).


et al. (2004) identified progressive greigite growth within the zone oflow magnetic mineral concentration at depths of 3–30 m below thesediment–water interface in the Korea Strait. Rowan et al. (2009)recognised similar features in sediments from the Oman and Californiamargins and reassessed records from the Ontong-Java Plateau(Tarduno, 1994), Japan Sea (Yamazaki et al., 2003), Argentine margin(Garming et al., 2005), and Niger Fan (Dillon and Bleil, 2006) to developa picture of down-coremagnetic property variations in a ‘Day plot’ (Dayet al., 1977) (Fig. 15a, b). Rowan et al. (2009) inferred progressivegreigite growth, starting with significant SP greigite concentrationsthat grew through the stable SD blocking volume with depth. Suchongoing greigite growth through early diagenesis will result in paleo-magnetic smoothing over periods of tens to hundreds of kiloyears, andhas profound implications for paleomagnetic records carried by greigite.Amodel for such greigite growth from Rowan et al. (2009) is illustratedin Fig. 15c, d.

The findings of Rowan et al. (2009) illustrate the care needed wheninterpreting paleomagnetic records that involve greigite. Conventionaltests that are used to demonstrate the antiquity of paleomagnetic sig-nalswill often give a false positive in such instances, including the rever-sals and fold tests. Artificial polarity sequences due to variable timing ofgreigite growth have been observed (Horng et al., 1998; Rowan andRoberts, 2005; Sagnotti et al., 2005a; Lucifora et al., 2012; Aben et al.,2014); the reversals test only failed in the case of Rowan and Roberts(2005) because large-scale tectonic rotations occurred between acquisi-tion of normal and reversed polarity magnetisations. Rowan andRoberts (2008) argued that fold tests are essential in tectonic studiesfor establishing the antiquity of magnetisations carried by greigite.However, folding could have occurredmillions of years after deposition,so that passing a fold test does not preclude acquisition of paleomagnet-ic signals over prolonged periods (cf. Liu et al., 2004; Rowanet al., 2009).

The above treatment highlights the importance of understandingthe challenges faced when diagenesis has had a pervasive, and poten-tially deceptively complex, influence on sedimentary magnetic records.A major development that would aid interpretation of magnetisationscarried by greigite would be to demonstrate that a signal is carried by

biogenic rather than diagenetic greigite. Diagenetic greigite can acquirea magnetisation at any time during diagenesis when dissolved iron andsulphide are present (Roberts andWeaver, 2005), but biogenic greigiteis expected to acquire a magnetisation at about the time of death of thehost magnetotactic bacterium. This should be within a few kyr of depo-sition if, in the worst-case scenario, the bacterium lived within the sed-iment. As indicated by Kopp and Kirschvink (2008), reports of putativegreigite magnetofossils are still relatively rare (Pósfai et al., 2001;Vasiliev et al., 2008); however, a recent report of ancient greigitemagnetofossils (Chang et al., 2014) suggests that it is possible to mag-netically detect such signals, which should simplify paleomagnetic in-terpretation of greigite-bearing sediments.

4.3.3. Pelagic marine environmentsPelagic marine sediments are deposited far from continents at water

depths of ~3000–6000m. In pelagic environments, organic matter con-tents are usually b0.2% in areas with low primary productivity (Morse,2005). Slowly deposited pelagic sediments (≤1 cm/kyr) tend to undergooxic diagenesis because organic carbon contents are low and dissolvedoxygen from theoverlyingwater column readily diffuses into sedimentsso that organic matter respiration occurs under oxic conditions(Table 2). Magnetic mineral diagenesis, therefore, is not as importantas in continental margin sediments, where organic carbon fluxes arehigh (Berner, 1980). As a result, oxic pelagic sediments often preservehigh-quality paleomagnetic recordswith little obvious indication of dia-genetic alteration. The spectrum of diagenetic conditions expected inpelagic marine environments is illustrated in Fig. 16 (Roberts et al.,2013a).

Pelagic red clays from the ultra-oligotrophic South Pacific gyre,where sedimentation rates are only ~1 mm/kyr (D'Hondt et al., 2009),provide an extreme illustration where oxic diagenesis occurs through-out the entire sedimentary succession (Fig. 16a). In this setting, surfacesediments have low organic carbon contents (0.2%) and dissolved oxy-gen has either beenmodelled or measured to have diffused through theentire sediment column to oxidise the underlying ocean crust basalt(Fischer et al., 2009; Expedition 329 Scientists, 2011). Hemi-pelagic

Fig. 14. Diagenetic modification of magnetic properties with depth in marine sediment cores from the southern Sea of Okhotsk. Water depths of the cores range from 778 m to 1443 m.χARM is the susceptibility of anhysteretic remanentmagnetisation (ameasure offine (submicron) ferrimagneticmineral concentration). χ and the colour code are as in Fig. 13. GHXX-XXXXis the name of the sediment core. Modified from Kawamura et al. (2007).


sediments from Eirik Drift, North Atlantic Ocean, provide an intermedi-ate example where oxic and ferruginous diagenesis persist over sedi-ment thicknesses of tens of metres (Fig. 16b; Kawamura et al., 2012).Although bottom waters are oxic, moderate rates of organic carbonburial result in sulphate reduction at depths of ~70 m below seafloor

(mbsf). However, the paleomagnetic signal is preserved to down-coredepths of 180mbsf (Kawamura et al., 2012) because of high initialmag-netite contents. Finally, in more productive waters, organic carbonfluxes can bemoderately high so that the oxic, nitrogenous,manganous,and ferruginous diagenetic zones all occur at shallow depths. In the


most productive waters, sulphidic diagenesis can control sedimentmagnetic properties, although these conditions occur at greater depthsthan in continental margin sediments (Tarduno, 1994; Fig. 16c).

Overall, diagenesis has variable effects on themagnetic properties ofpelagic sediments, ranging fromminimal to controlling influences. Oxicto nitrogenous conditions are ideal for iron oxide preservation, whichoften makes pelagic sediments superb recorders of paleomagnetic in-formation (e.g., Opdyke and Henry, 1969; Lowrie et al., 1980; Channellet al., 2003; Lanci et al., 2004; Florindo and Roberts, 2005; Channellet al., 2013; Roberts et al., 2013a). Nevertheless, prolonged oxygen dif-fusion into the sediment can lead to partial or complete oxidation of de-trital magnetite or titanomagnetite to form maghemite (e.g., Cui et al.,1994) or titanomaghemite (e.g., Xuan and Channell, 2010) rims on thesurface of (titano)magnetite grains. Complexities resulting from suchoxidation canmake oxic pelagic sediments unsuitable for paleomagnet-ic recording (e.g., Xuan and Channell, 2010). Haematite or goethite for-mation in oxic diagenetic environments can also be important in somepelagic and hemi-pelagic settings (e.g., Henshaw and Merrill, 1980;Channell et al., 1982; Bouilloux et al., 2013), while diagenetic magnetitedissolution and destruction of the paleomagnetic record can be impor-tant below productive surface waters. Detrital iron oxide dissolutionhas been documented in biologically productive deep marine settings,including the Tyrrhenian Sea (Channell and Hawthorne, 1990), the Ara-bian Sea (Bloemendal et al., 1992, 1993; Rowan et al., 2009; Bouillouxet al., 2013), the South Atlantic Ocean (Bloemendal et al., 1992;Channell and Stoner, 2002), the North Atlantic Ocean (Robinson andSahota, 2000; Robinson et al., 2000; Kawamura et al., 2012), and theOntong-Java Plateau (Tarduno, 1994, 1995; Tarduno and Wilkison,1996; Tarduno et al., 1998).

While it is argued above that oxic to nitrogenous conditions can beideal for preserving paleomagnetic signals, complex magnetic signalsare recorded by red clay sediments that are deposited in pelagic settingsthat lie below the calcite compensation depth in deep ocean basins farfrom land. Development of paleomagnetic instability at depths of afew metres below the seafloor has been reported widely for thesesediments (e.g., Opdyke and Foster, 1970; Kent and Lowrie, 1974;Johnson et al., 1975; Henshaw and Merrill, 1980; Yamazaki andKatsura, 1990; Yamazaki and Ioka, 1997). Themagnetic instability of pe-lagic red clays has been attributed to low-temperature oxidation of sta-ble SD magnetite, which is now known to be carried by magnetitemagnetofossils (Yamazaki and Shimono, 2013), so that it forms smallermagnetically unstable SP maghemite particles (Kent and Lowrie, 1974;Johnson et al., 1975). Such oxidation would be expected with down-ward diffusion of oxygen into the sediment, as described above inSection 4.2.1.1. The low sedimentation rates of pelagic red clays meanthat they are not widely used for paleomagnetic studies, but they docarry important records of eolian dust deposition (Yamazaki and Ioka,1997). Magnetic polarity stratigraphy provides an important datingtool for such sediments, which makes it valuable to better understandthe mechanisms by which pelagic red clays acquire an unstablemagnetisation with depth.

4.3.4. Lake sedimentsLake sediments generally contain plentiful sources of organicmatter

that drive magnetic mineral diagenesis. However, the sulphate contentof freshwater (between 100 and 200 μM kg−1) is much lower than inseawater (28 mM kg−1 at a salinity of 35 permil). As a result, manga-nese and iron reduction are generally major processes in lake sedi-ments, while sulphate reduction is expected to be less important.Methanogenesis is the dominant terminal diagenetic process in fresh-water settings (Capone and Kiene, 1988; Nealson, 1997) compared tosulphate reduction in marine sediments (although the importance ofsulphate reduction in marine sediments has recently been proposed tobe less important than was traditionally thought; Bowles et al., 2014).A key consequence of the reduced relative importance of sulphidic dia-genesis in lake sediments is that detrital magnetic minerals are morelikely to survive diagenesis than in sulphidic marine sediments. Never-theless, sulphate values are not zero and magnetite dissolution is awidely reported process in organic-rich lake sediments (e.g., Andersonand Rippey, 1988; Snowball, 1993a, 1993b; Thouveny et al., 1994;Rosenbaum et al., 1996; Snowball, 1996; Stockhausen and Zolitschka,1999; Demory et al., 2005; Ao et al., 2010; Nowaczyk, 2011; Su et al.,2013; Fu et al., 2015). Much iron oxide dissolution in lake sedimentscould occur via iron reduction; nevertheless, widespread iron sulphideformation indicates that sulphidic diagenesis is not a negligible processin lake sediments. For example, greigite formation and preservation hasbeenwidely reported in lake sediments (Skinner et al., 1964; Dell, 1972;Giovanoli, 1979; Snowball and Thompson, 1988, 1990a, 1990b, 1992;Hilton, 1990; Snowball, 1991, 1996; Roberts et al., 1996; Turner, 1997;Hu et al., 1999; Peters and Turner, 1999; Reynolds et al., 1999;Stockhausen and Zolitschka, 1999; Demory et al., 2005; Frank et al.,2007a, 2007b; Ron et al., 2007; Nowaczyk, 2011). These observationsdemonstrate that sulphidic diagenesis cannot be ignored in lake sedi-ments. Lakes are dynamic systems in which lakewater sulphate and or-ganic carbon contents can vary significantly through time. This makes itreasonable to consider lake sediment diagenesis as a non-steady stateprocess (see Section 4.4).

Continental weathering is the principal source of sulphate to theoceans. This is also the case for lakes, with drainage of soils containinggypsum,weathering of rocks containing pyrite, or decomposition of vol-canic ash providing major sources of sulphate. The smaller volume oflakes compared to the oceanmakes themmore immediately susceptibleto anthropogenic nutrient inputs from effluent, including agriculturaland industrial runoff; modern europhication can have important conse-quences for magnetic mineral preservation in lake sediments(e.g., Anderson and Rippey, 1988; Hilton, 1990; Ariztegui and Dobson,1996). Also, sulphate can accumulate in saline lakes where the rate ofevaporation or withdrawal of freshwater exceeds the rate of replace-ment. Salt concentrations can exceed those of seawater, at whichpoint the lake is referred to as hypersaline. Many lakes in arid zones,therefore, provide a record of varying freshness/salinity in relation tothe balance between precipitation/inflow and evaporation. Changingfreshness, which affects the sulphate content of lakewater, and variable

Fig. 15. Illustration of progressive greigite formation with depth in hemi-pelagic marine sediments. (a) Day plot (after Day et al., 1977) of hysteresis ratios (Mr = saturation remanentmagnetisation; Ms = saturation magnetisation; Bcr = coercivity of remanence; Bc = coercive force) for sediments from Oman margin core CD143-55705, in which progressive down-corevariations shift the data distribution to higher Bcr/Bc values in Zone 2 (shaded grey in the plot of down-core natural remanent magnetisation (NRM) variations on the right-hand side),followed by a return tomore single domain (SD)-like values in the lowest interval of the core (Zone 3). The association of these changes with progressive pyritisation and greigite forma-tion suggests that the observed looping is due to progressive superparamagnetic (SP) greigite formation in Zone 2,which then grew through the stable SD blocking volume in Zone 3. Thetheoretical SD–MDmixing curve for (titano-)magnetite is fromDunlop (2002), and the SD–SPmixing trend represents the uppermost boundary of the envelope of SD–SPmixtures fromgreigite-bearing sediments inNewZealand (RowanandRoberts, 2006). The SD, pseudosingle domain (PSD) andmulti-domain (MD)fields for (titano-)magnetite are plotted for referenceonly. (b) Day plot of progressive down-core looping of hysteresis parameters associated with sulphidic diagenesis in marine sediments from the Oman and northern California margins(Rowan et al., 2009), Korea Strait (Liu et al., 2004), Ontong-Java Plateau (Tarduno, 1994, 1995), Niger Fan (Dillon and Bleil, 2006), Oregon margin (Karlin, 1990a), Japan Sea (Yamazakiet al., 2003), and Argentine margin (Garming et al., 2005). (c) Simplifiedmodel for progressive changes in magnetic mineral content of sulphidic sediments, illustrating progressive dis-solution of SD and PSDmagnetite (red area/line) and formation of SP/SD greigite (grey line/area) at the sulphate-methane transition (SMT), the delay before greigite grows to SD size, andcontinuous greigite growth in Zone 3. High coercivity minerals (e.g., haematite) are excluded for clarity, but are expected to follow the progressive down-core dissolution trend describedby Liu et al. (2004). (d)Modifiedmodel inwhich a recent upwardmigration of the SMT in the sediment column is incorporated, leaving an interval enriched in PSD/MDmagnetite (lowerZone 2) between the old and new dissolution fronts. All sub-plots were modified from Rowan et al. (2009).


organic carbon content can, therefore, provide records of non-steadystate diagenesis that reflect environmental variations (e.g., Snowballand Thompson, 1990a; Roberts et al., 1996; Peck et al., 2004; Ao et al.,2010). Such considerations are also relevant to marginal seas, such asthe Baltic Sea, Black Sea or Caspian Sea, that have been isolated fromthe open ocean for significant parts of their history. Transitions to brack-ish conditions particularly favour greigite formation and preservation(e.g., Sohlenius, 1996; Jelinowska et al., 1998; Strechie et al., 2002;Jørgensen et al., 2004; Neretin et al., 2004; Chang et al., 2014;Holmkvist et al., 2014). Influxes or removal of seawater into/from coast-al lakes through sea level variations, post-glacial rebound/glacial subsi-dence, or tsunami inputs can give rise to further special cases of non-steady state diagenesis. Sulphidic diagenesis, including dissolution ofmagnetite and authigenic greigite formation, is much more likely todominate when seawater sulphate becomes available in lacustrine sed-iments (e.g., Snowball and Thompson, 1990a; Wang et al., 1999). Over-all, the major forms of diagenetic magnetic mineral modificationobserved in marine sediments also occur in lake sediments. Non-steady state diagenesis is likely to dominate, so that diagenetically

influenced magnetic properties can provide useful paleoenvironmentalinformation in lake sediments (e.g., Thouveny et al., 1994; Roberts et al.,1996; Peck et al., 2004; Ao et al., 2010).

The fact that methanogenesis is the dominant terminal diageneticprocess in freshwater settings (Capone and Kiene, 1988; Nealson,1997)means that the processes discussed above in relation tomethanicmagnetic mineral diagenesis are likely to be important in lake sedi-ments. For example, Sivan et al. (2011) argued that iron reduction inthe methanic zone can drive AOM in lake sediments. Thus, magneticiron sulphide formation via AOM could be important for consideringmagnetic mineral diagenesis in lake sediments. A key issue then be-comes the depth at which AOM occurred, and, therefore, the offset be-tween the age of the recorded magnetic signal and that of the hostsediment. Sivan et al. (2011) documented iron-mediated AOM at shal-low depths (20–30 cm below the sediment–water interface) in LakeKinneret (Sea of Galilee), Israel, which would, if it gave rise to greigiteformation, only cause a mild delay in magnetic signal acquisition. Nev-ertheless, if AOM occurred deeper in the sediment column, then thepossibility of late remagnetisation, as has been widely observed in

Fig. 16. Schematic illustration of sedimentary pore water profiles associatedwith organic matter diagenesis in three pelagic settings (modified from Roberts et al. (2013a)), with progres-sively increasing influence of organic matter diagenesis to illustrate the spectrum of diagenetic environments in pelagic settings. Depths are reported in metres below seafloor (mbsf).(a) Pelagic red clay sediments from the ultra-oligotrophic South Pacific gyre, where sedimentation rates are only ~1 mm/kyr (D'Hondt et al., 2009). Non-zero dissolved oxygen concen-trations have beenmeasured through the entire sediment column at IODP Site U1365 (Expedition 329 Scientists, 2011). Any organicmatter delivered to the seafloor is rapidly oxidised, sothat the sediments only experienced oxic diagenesis with oxygen diffusing through the entire sediment column to oxidise the underlying oceanic lithosphere basalt. (b) Hemi-pelagicsediments from IODP Site U1305, Eirik Drift, North Atlantic Ocean (after Kawamura et al., 2012). Bottom waters are oxic, but moderate rates of organic carbon burial result in sulphatereduction at depth. High initial magnetite contentsmean that the paleomagnetic signal is preserved to 180mbsf. Porewater Fe2+ data are noisy and are not shown. (c) Pelagic carbonatesfrom ODP Hole 806A, Ontong-Java Plateau (after Tarduno, 1994); the profile is shown only for the uppermost 12 mbsf. The paleomagnetic signal is largely destroyed below this depth.


methanic marine sediments (e.g., Roberts and Weaver, 2005; Sagnottiet al., 2005a; Larrasoaña et al., 2007; van Dongen et al., 2007), willneed to be considered.

4.3.5. SoilsThe discussion above is explicitly related to sediments rather

than soils. However, organic carbon degradation also occurs in soilenvironments, and the framework outlined above in relation tomag-netic mineral diagenesis and the energetics of organic carbon degra-dation through microbial metabolism from oxic to anaerobic

respiration are equally applicable (Fig. 17). Pedogenesis adds com-plexity to the above-described perspective on diagenesis. For exam-ple, manymagnetic mineral formation and transformation processeshave been proposed for oxic soils. It is beyond the scope of this paperto discuss such processes in detail; readers are referred to the follow-ing papers for outlines of magnetic mineral formation pathways insoils (e.g., Maher and Taylor, 1988; Zhou et al., 1990; Dearing et al.,1996; Maher, 1998; Torrent et al., 2006; Liu et al., 2007; Blundellet al., 2009; Boyle et al., 2010). For the purposes of this paper, it issufficient to say that pedogenic formation of fine-grained iron

Fig. 17. Schematic cartoonof non-steady state diagenetic processes that vary in time and space in terrestrial sedimentary environments,with redox zonations associatedwith groundwaterflow, reactive transport of inorganic contaminants and wetlands. Insets are illustrations of iron mineral precipitation in different diagenetic zones (modified from Davis et al. (2004)).

Fig. 18. Photographic illustrations of iron diagenesis in soils. (a) A gley podzol fromOkarito,West Coast, South Island, New Zealand. The base of the profile consists of Last Glacial outwashgravel that is cemented with iron and humus. The white E horizon occurs in loess and the upper part of the profile is peat. The scale graduations on the tape (5 cm each) indicate that thehole is ~1.3m deep. Note the present position of thewater table at the bottom of the soil pit. Podzols form in environments with high precipitation. Leaching of iron and aluminiumout ofthe E horizon has led to formation of an iron pan that has impeded drainage. Podzols are common in the UK, Europe, China, North America, southern Australia and New Zealand.(b) Holocene gleyed Kairanga silty clay loam (b2000 years old) from near Palmerston North, North Island, New Zealand. Waterlogging has led to reduction of Fe3+ to Fe2+, as indicatedby grey colouration below a depth of 25 cm. Lowering of the water table and reoxygenation has led to formation of yellowmottles within the gleyed clay. Note the present position of thewater table at the bottom of the soil pit. The images in (a) and (b) are fromMolloy (1998) and are used with permission of the New Zealand Soil Science Society. (c) Nodular haematitemottles in upper Lower Pleistocene sandy clays of the Ochre Cove Formation at Sellicks Beach, south of Adelaide, Australia. Pillans and Bourman (2001) inferred that the haematitemottlesformed by subsurface weathering associated with a fluctuating water table below the active soil-forming zone under a higher rainfall regime than the present climate.


(oxyhydr-)oxides gives rise to magnetic mineral enhancement inmany topsoils. In discussing soil diagenesis, it is the diagenetic trans-formation of detrital and pedogenic magnetic minerals that is of in-terest rather than pedogenesis, although repeated wetting-dryingcycles that are relevant to diagenesis can be an important part ofpedogenic magnetic mineral formation and transformation. Thismakes it fundamentally difficult to disentangle the associated pro-cesses; this reality should be borne in mind when considering soildiagenesis.

Perpetually water logged soils provide the simplest case ofmagneticmineral diagenesis in soils, where reductive diagenesis is analogous tothat in lake sediments (Fig. 17). Magnetite dissolution is widespreadin such gleyed soils (Fig. 18a) (e.g., Dearing et al., 1996; Maher, 1998;Guo et al., 2001; Chlachula, 2003; Blundell et al., 2009). Redox-relatedtransformations of iron-bearing minerals in soils outside of wetlandsare more complicated and are better considered in terms of non-steady state diagenesis (see Section 4.4). Soils undergo major non-steady state changes with repeated wetting and drying cycles (due torainfall, irrigation or groundwater variations). Diffusive reintroductionof oxygen during drying produces localised redox microgradientsaround soil particles, cracks, pores and root channels, which adds tocomplexities due to variable temperature, nutrient input and/oraddition of dissolved gases from plant roots, and bioturbation(e.g., Nealson, 1997). Localised mottling (Fig. 18b, c) is common in soilprofiles and is a manifestation of imperfect drainage, with repeatedwetting-drying cycles (e.g., Simonson and Boersma, 1972; Venemanet al., 1976; Duchafour, 1982). In mottled soils, oxidative precipitationof Fe3+ minerals occurs during dry periods, reductive dissolution andFe3+ migration (often over short distances) occurs during wet periods,leaving Fe2+ in the reduced water logged soil, and re-precipitation ofFe3+ mineral phases occurs during subsequent drying. Such redox cy-cling can occur repeatedly and gives rise to blotches of grey andreddish-brown-yellow colours at the same depth in a soil profile(Fig. 18b, c). Soil colour is a direct result of these redox reactions. Redis common in well-aerated soils (rich in Fe3+; Fig. 18c), yellow is com-mon in soils with intermediate aeration (Fig. 18b) and blue and grey areindicative of poor aeration (rich in Fe2+).Microbialmanganese and ironreduction in soils can give rise to precipitation of manganese and ironnodules in soils, which occur commonly alongside soil mottles(e.g., Simonson and Boersma, 1972; Veneman et al., 1976; Duchafour,1982; Kraus and Aslan, 1993).

Dissimilatory iron-reducingbacteria are strict anaerobes (Lovley andPhillips, 1986; Lovley et al., 1987), so they needwater logged soil condi-tions to reduce Fe3+ to Fe2+.While dissimilatory iron-reducing bacteriahave been argued to be unnecessary for pedogenic magnetic mineralformation (Maher and Taylor, 1988), they are widespread in waterlogged soils (Lovley, 1995) and their importance for iron cyclingin such environments makes them a key potential source ofnanoparticulate soil magnetite. Uncertainty about the abundance of dis-similatory iron-reducing bacteria and their role in pedogenic magneticparticle formation during wetting-drying cycles in soils limits our un-derstanding of magnetic mineral enhancement in soils (Boyle et al.,2010).

Few discussions of diagenesis consider soils. The present treatmentis brief because major processes that are also applicable to soils arediscussed in detail above. Key issues relate to complexities due to thenon-steady state nature of soil diagenesis and the challenge of discrim-inating between pedogenic and diagenetic processes when investigat-ing paleosols (e.g., Kraus and Aslan, 1993). Soils are a special case ofnon-steady state diagenesis; the following discussion of non-steadystate diagenesis focuses on aquatic sedimentary environments.

4.4. Non-steady state diagenesis

Much of the above discussion assumes that diagenesis has occurredunder steady state conditionswhere pore waters are in equilibrium and

the pore water zonation does not change significantly through time.This condition is most likely to be satisfied in pelagic environmentswhere generally loworganic carbonfluxes are often approximately con-stant so that steady state diagenesis occurs. However, biological produc-tivity can vary significantly through time as a result of intermittentsupply of key limiting micronutrients required by plankton (e.g., ironfertilisation of oligotrophic waters through supply of eolian dust;Martin et al., 1991). Time varying pulses of organic carbon delivery tothe seafloor or lake floor can then give rise to marked changes in mag-netic mineral diagenesis. Such non-steady state diagenesis can be driv-en by orbitally-controlled productivity cycles that result in varyingbacterial magnetite preservation (e.g., Hesse, 1994; Lean and McCave,1998) or magnetite dissolution cycles (e.g., Tarduno, 1992; Dekkerset al., 1994; van Os et al., 1994; van Santvoort et al., 1997; Larrasoañaet al., 2003a, 2006; Drab et al., 2015). Instantaneous delivery of en-hanced organic carbon concentrations to the seafloor by turbidites orother sediment mass flow events (e.g., Robinson et al., 2000; Hensenet al., 2003), sea-level induced pore fluid variations in continental mar-gin sediments (Fig. 19a; e.g., Oda and Torii, 2004), or sea level controlledsedimentation rate variations (Riedinger et al., 2005, 2014; Fu et al.,2008; März et al., 2008; Abrajevitch and Kodama, 2011) can also enablevariable formation and/or preservation of magnetic minerals. Glacial-interglacial changes in oceanic upwelling can change nutrient condi-tions that cause the position and thickness of an oxygen minimumzone to vary in the water column, which gives rise to changes in irondiagenesis at the seafloor (Fig. 19b; e.g., Scholz et al., 2014). Non-steady state magnetite dissolution associated with relatively short-lived enhanced organic carbon burial events can explain cyclicity be-tween relatively strong and weak magnetisations in some instances(e.g., Itambi et al., 2010). Alternatively, enhanced organic carbon supplycan result in ferruginous diagenetic conditions that liberate Fe2+ intosedimentary porewaters, thereby providing the iron needed to enhancemagnetite production by magnetotactic bacteria (e.g., Roberts et al.,2011a; Larrasoaña et al., 2012; Yamazaki and Ikehara, 2012). In this sce-nario, enhanced organic carbon supply can cause increasedmagnetisations. Discrimination between such scenarios requires use ofgeochemical and magnetic proxies for diagenesis. Readers are referredto Kasten et al. (2003) for a detailed treatment of non-steady state dia-genetic processes and signals. Non-steady state diagenesis is more im-portant than steady state diagenesis in many settings, particularly inlake sediments, but also in many marine environments. Signatures ofnon-steady state diagenesis can provide valuable insights into the cli-matic drivers of diagenetic variations (see examples below), althoughnon-steady state diagenesis can also obscure climate-related magneticproxies (e.g., Abrajevitch et al., 2009).

Deposition of organic-rich layers, known as sapropels, in the EasternMediterranean Sea provides a clear example of non-steady state mag-netic mineral diagenesis in deep marine environments (Dekkers et al.,1994; van Os et al., 1994; van Santvoort et al., 1997; Passier et al.,2001; Passier and Dekkers, 2002; Larrasoaña et al., 2003a, 2006; Liuet al., 2012b). Normal oxic pelagic sedimentation in the Eastern Medi-terranean has often been disrupted during orbital precession-driven in-solation maxima (e.g., Hilgen, 1991), when the African monsoonintensified and penetrated northward into the Mediterranean catch-ment (Rossignol-Strick, 1983; Lourens et al., 2001; Larrasoaña et al.,2003b). Coupled with wetter conditions in Europe (e.g., Rohling andHilgen, 1991), increased freshwater discharge into the EasternMediter-ranean increased the buoyancy of surface waters that led to deteriora-tion of deep-water ventilation and enhanced export of organic carbonto the seafloor that produced conditions conducive to sapropel forma-tion (e.g., Rohling, 1994; Rohling et al., 2015).

These pulses of enhanced organic carbon deposition and preserva-tion at the seafloor induced non-steady state diagenetic changes thatare reflected in the sediment magnetic record (Larrasoaña et al.,2003a, 2006; Liu et al., 2012b). Typical variations are illustrated in a car-toon in Fig. 20. During times of sapropel formation, Eastern


Mediterranean bottom waters were sulphidic, and enhanced organiccarbon deposition gave rise to downward sulphidisation belowsapropels. This caused dissolution of ferrimagnetic iron (oxyhydr-)ox-ides, pyrite formation, and loss of magnetisation in the sedimentthrough which the downward-moving sulphidisation front passed.When periods of sapropel formation ended, conditions again becameoxic,which allowed oxygen diffusion into the sediment. Oxidation of re-duced iron in sapropels led to formation of iron (oxyhydr-)oxides, in-cluding magnetite, which enhanced magnetisations at the oxidation

front at the top of the sapropel (Passier et al., 2001; Passier andDekkers, 2002; Larrasoaña et al., 2003a, 2006; Liu et al., 2012b). Back-ground sediments deposited under oxic conditions, which include sig-nificant Saharan dust contents (Larrasoaña et al., 2003b), havemagnetisations that are intermediate between theweakmagnetisationsof the sulphidised sediments below and within sapropels, and thestrong magnetisations in the oxidation front at the top of sapropels.These non-steady state diagenetic variations produce the characteristicmagnetic properties shown in Fig. 20.

Fig. 19. Schematic cartoon of non-steady state diagenetic processes that vary in time and space in coastal and continental margin sedimentary environments. (a) Representation of redoxand freshwater-saltwater interface changes in the coastal zone associated with sea level variability on glacial-interglacial timescales (modified fromMulligan and Charette (2009)). Sub-marine discharge of fresh groundwater is an important process in coastal and continental shelf environments (Mulligan and Charette, 2009; Post et al., 2013) andmovement of the fresh-water-saltwater interface could contribute to non-steady state diagenesis for considerable distances along continentalmargins. (b) Representation of sedimentary diagenetic zonations ona continental margin in relation to glacial-interglacial variability in the position of an oxygen minimum zone (OMZ) related to variable upwelling of nutrient enriched waters (modifiedfrom Scholz et al. (2014)). Surface sediments can become sulphidicwhere theOMZ is strong (and bottomwater ventilation isweak). Ferruginous diagenesis in surface sediments providesa source of iron to the ocean, whereas sulphidic conditions cause iron retention within iron sulphides. Also shown is a global map of ocean oxygen content at a water depth of 200 m;oxygen minima in the water column are usually associated with zones of nutrient upwelling. Data are from the NOAAWorld Ocean Atlas.


Magnetic property variations illustrated in Fig. 20, along with strati-graphic variations in Ti/Al and haematite contents that indicate Saharandust influxes to the Eastern Mediterranean (Lourens et al., 2001;Larrasoaña et al., 2003b), provide a climatically (summer insolation)driven magnetic signal that responds directly to orbital forcing(Fig. 21). Eastern Mediterranean sapropels occur in clusters associatedwith orbital eccentricity maxima and precession minima (Hilgen,1991). The types of magnetic property variations shown in Fig. 20 arerepresentative of a ‘type 2’ sapropel (Larrasoaña et al., 2003a). In con-trast, during eccentricity minima, sapropel preservation is much lessmarked. Paleoproductivity proxies, such as Ba/Al, reveal that elevatedorganic carbon concentrations were deposited on the seafloor, butwere not preserved (arrows in Fig. 21). Such missing sapropels escapevisual identification due to complete removal by post-depositional oxi-dation, but can be detected geochemically (e.g., Higgs et al., 1994; vanSantvoort et al., 1997; Wehausen and Brumsack, 2000; Calvert andFontugne, 2001). These ‘type 3’ sapropels lack a magnetite dissolutionfront below the sapropels, but can be identified magnetically by asmall peak produced by an oxidation front at the top of the missingsapropel (Fig. 22). Finally, during high-amplitude eccentricity maxima,another type of sapropel signature is observed (‘type 1’). Sedimentsbetween sapropels no longer return to yellow–brown–red colours, butremain blue-grey (Fig. 23). These sapropels are enriched in redox-sensitive trace metals (Mo, V) that are sensitive indicators of bottomwater anoxia (e.g., Morford and Emerson, 1999; Algeo and Lyons,2006). Magnetisations are reduced to near-zero values (Fig. 24),which reflect pervasive sulphidisation and dissolution of magnetite insapropels and intervals between sapropels. Some type 1 sapropels

are magnetically enhanced due to magnetic iron sulphide formation(Roberts et al., 1999).

Observed differences in sapropel magnetic properties provide auseful means of determining past variations in Eastern Mediterra-nean deep-water ventilation (Fig. 23) (Larrasoaña et al., 2003a).The different sapropel types are evident in a 2.2-million-year recordof anhysteretic remanent magnetisation (ARM) variations. The vari-ations are consistent with sediment colour variations, which arestrongly modulated by orbital variations (Fig. 23). These resultsdemonstrate strikingly how diagenetically controlled magneticproperty variations can be used to obtain valuable paleo-proxy re-cords and to make inferences about the climatic driving forces ofnon-steady state diagenesis. In contrast to pervasive sulphidic dia-genesis, which destroys the detrital magnetic signal that is usuallythe object of paleomagnetic and environmental magnetic analysis,non-steady state diagenesis can provide extremely useful cluesabout a range of climatic processes (e.g., Larrasoaña et al., 2003a;Blanchet et al., 2009; Nowaczyk et al., 2012).

Sedimentation and the supply of reactive organic matter and iron-bearing minerals is non-uniform in many environments, and porewater zonations can be perturbed by instantaneous changes due, for ex-ample, to seismicity, gravity flows, etc., and gradual changes due, for ex-ample, to sedimentation rate, lake level or sea level changes. Suchchangesmean that non-steady state diagenesis is fundamentally impor-tant and should not be ignored. The importance of non-steady state dia-genesis means that it is described in various places throughout thispaper in relation to numerous depositional settings, which leads to lim-ited but unavoidable repetition.

Fig. 20. Illustration of the effects of variable bottomwater ventilation onnon-steady state diagenesis in EasternMediterranean sediments. During times of sapropel (organic-rich sediment)formation, bottom waters were sulphidic so that downward diffusion of H2S sulphidises sediments below sapropels, which gives rise to magnetite dissolution, pyrite formation, loss ofmagnetisation, and imparts a grey colour to the sediments. When bottom waters are reventilated, oxygen diffuses downward into the formerly sulphidic sediments, which producesan oxidation front where authigenic magnetic minerals form (Passier et al., 2001; Passier and Dekkers, 2002). This sequence of bottomwater ventilation variations gives rise to the illus-trated magnetic property variations (characteristic of a type 2 sapropel; Larrasoaña et al., 2003a).


4.5. Diagenetic microenvironments

Localised iron sulphide occurrences are observed commonly in sed-imentary environments that appear to have otherwise remained oxicduring and after deposition (Fig. 24). Such sulphidisation occurs aroundisolated pieces of decaying organic matter in environments where theoverall supply of organic matter is insufficient to drive sulphidisationof the entire sediment volume. If dissolved iron is available within sed-imentary pore waters, due to active iron reduction, it will only react toform iron sulphides in microenvironments with available dissolved sul-phide. H2S that diffuses out of decaying organic matter will react withavailable dissolved iron and will be sequestered as iron sulphide(Figs. 11f and 24) that precipitates, and is confined around, the organicmatter (Allen, 2002). Diagenetic microenvironments are more impor-tant in pelagic sediments than in continental margin sediments where

reducing diagenetic conditions are more pervasive. For example,reducing microenvironments can develop within microfossil tests(e.g., Stumm and Morgan, 1996), faecal pellets (e.g., Demory et al.,2005) or around larger pieces of organicmatter, where sedimentary py-rite formation is only localised (Fig. 24). Such situations are evident bypreservation of iron oxides in unaffected parts of the sediment and bypyrite or greigite formation, with lack of detrital iron oxides, in locallyreduced sediments (e.g., Demory et al., 2005). Iron sulphide nodules,which range considerably in size (Figs. 8 and 11), are also common inpervasively reduced sediments (e.g., Florindo and Sagnotti, 1995; Jianget al., 2001; van Dongen et al., 2007), again because of sulphidisationof macro-scale organic matter (Figs. 8 and 11f) (Grimes et al., 2001;Rickard et al., 2007) or of micro-scale organic matter within microfossiltests (Fig. 9) (Roberts and Turner, 1993; Roberts et al., 2005). Iron sul-phide nodules often contain ferrimagnetic greigite (Fig. 8c, g, h; Jiang

Fig. 21. Illustration of magnetic property variations for type 2 and type 3 sapropels due to non-steady state diagenesis in Eastern Mediterranean sediments. See text for a detailed expla-nation of sapropel types. Ba/Al provides a measure of paleoproductivity and Ti/Al provides a measure of Saharan dust input (fromWehausen and Brumsack, 2000). HIRM variations pro-vide a measure of haematite inputs, which are consistent with Ti/Al variations, and indicate variations in Saharan dust input due to African monsoon variations (Larrasoaña et al., 2003b),which responds directly to orbital forcing, as indicated by the correlation with summer insolation at 65°N (from Laskar, 1990). Insolation maxima correspond to precession minima. Or-ganic carbon deposition during periods of normal sapropel formation gives rise to magnetic signatures associatedwith type 2 sapropels (Fig. 20). Post-depositional oxidation gives rise tooxidation and removal of visible sapropels, which are classified as type3 sapropels (indicated by arrows). Data are fromOceanDrilling Program(ODP) Site 967. Depths are from the revisedcomposite metres depth (rmcd) scale of Emeis et al. (2000). Modified from Larrasoaña et al. (2003a).


et al., 2001; Neretin et al., 2004; Roberts andWeaver, 2005; vanDongenet al., 2007; Florindo et al., 2007), monoclinic pyrrhotite (Fig. 12d-k;Dinarès-Turell and Dekkers, 1999; Weaver et al., 2002; Larrasoañaet al., 2007; Roberts et al., 2010), and antiferromagnetic hexagonal pyr-rhotite (Fig. 8i-l; van Dongen et al., 2007), as discussed above.

4.6. Silica diagenesis and magnetic minerals

Stratigraphic variations of pore water and magnetic property varia-tions and thermodynamic calculations were used by Florindo et al.(2003) to suggest thatmagnetite is unstable in thepresence of dissolvedsilica. Dissolved pore water silica can result from dissolution of the sili-ceous tests of diatoms and radiolaria. The organic remains of thesebiosiliceous organisms contribute to the organic carbon content of sed-iments, which makes it important to discriminate between magneticmineral instability due to silica diagenesis compared to conventionalorganic matter-related reductive diagenesis. Florindo et al. (2003) ob-served low magnetite concentrations in marine sediments with highdissolved silica concentrations, where pore water sulphate remained

at typical seawater values, which suggests that sulphate reduction hasnot occurred. However, past non-steady state organic carbon burialevents in sediments that currently do not support active sulphate re-duction is possible and would lead to false attribution to silica-relatedmagnetite diagenesis rather than to past sulphidic diagenesis. Florindoet al. (2003) assessed this possibility using sediments with low biogenicsilica productivity and low organic carbon contents, but with interstitialsilica due to dissolution of volcanic ash, and found that magnetisationsare weak. They, therefore, predicted that magnetite dissolution will bea common feature in siliceous sedimentary environments. Subsequentstudies have reported associations of low magnetite concentrationswith high biogenic silica concentrations (e.g., Wetter et al., 2007;Hillenbrand et al., 2010). While these observations might support theconclusions of Florindo et al. (2003), theymight also simply reflect con-ventional organic carbon diagenesis. Other studies have reported fea-tures such as authigenic smectite formation in zones with eitherelevated biogenic or volcanogenic silica contents (e.g., Demory et al.,2005; Homoky et al., 2011) that Florindo et al. (2003) predicted will re-sult from magnetite dissolution in non-sulphidic environments. In

Fig. 22. Illustration ofmagnetic property variations associatedwith type 1–3 sapropels due to non-steady state diagenesis in EasternMediterranean sediments. Sapropel types 2 and 3 areas in Fig. 21. Weak magnetisations (ARM) associated with type 1 sapropels are associated with pervasive magnetic mineral dissolution during periods of prolonged bottomwater anoxiaduring and between periods of sapropel formation. Parameters are as in Fig. 21; Mo/Al and V/Al are proxies for bottomwater ventilation. Data are from ODP Site 966 and rmcd depths arefrom Emeis et al. (2000). Modified from Larrasoaña et al. (2003a).


contrast, Abrajevitch et al. (2011) reported preservation of fine-grainedmagnetite in red cherts, which is not consistent with the proposal ofFlorindo et al. (2003).

Whether magnetite dissolution associated with silica diagenesis isimportant remains lesswell resolved than for sulphidic diagenetic envi-ronments. It has been proposed recently that silica precipitation iscatalysed by adsorption to freshly precipitated iron oxide surfaces(Meister et al., 2014). If iron diagenesis is important in chert formation,further possible relationships between iron and silica diagenesis andmagnetic minerals remain to be elucidated.

4.7. Relict magnetic mineral assemblages in sulphidic sediments

Ilmenite and Fe–Cr and Fe–Mn spinels are common relict minerals(Fig. 25) in diagenetically reduced sediments (e.g., Roberts and Turner,1993; Hounslow et al., 1995; Hounslow, 1996; Wilson and Roberts,1999; Garming et al., 2005; Itambi et al., 2010; Nowaczyk, 2011) due

to their resistance to reductive dissolution under sulphidic conditions(Canfield et al., 1992). Increasing cation substitution can stabilise suchminerals against sulphidic dissolution because substitution reducesthe concentration of the electron acceptor Fe3+.

It is worth exploring briefly the potential paleomagnetic importanceof detrital ilmenite particles in diagenetically reduced sediments. Ilmen-ite is commonly observed in magnetic mineral extracts from suchsediments (e.g., Roberts and Turner, 1993; Hounslow et al., 1995;Hounslow, 1996; Wilson and Roberts, 1999; Itambi et al., 2010;Nowaczyk, 2011; Chang et al., 2015), and its presence is often attributedto its paramagnetic response to the strong magnetic fields applied dur-ingmagneticmineral extraction. It is, therefore, usually ignored in termsof paleomagnetic signal recording. However, Robinson et al. (2002) pro-posed that lamellar magnetism provides a mechanism by which ilmen-ite can record magnetic signals. In this mechanism, pervasive formationof exsolution lamellae in slowly cooled igneous rocks is argued to pro-duce a ferrimagnetic substructure due to Fe2+–Fe3+ charge imbalance

Fig. 23. Illustration of the use ofmagneticmineral diagenesis in sapropels to infer long-termbottomwater ventilation variations in the EasternMediterranean Sea (using the framework inFigs. 20–22). Bottom: Sapropel distribution (vertical grey boxes) and anhysteretic remanent magnetisation (ARM) data for ODP Site 966 between 1.9 and 4.1Ma. Lowermiddle: Sapropelclassification, which indicates variations from perpetual bottomwater anoxia (type 1) to bottomwatter anoxia only during times of sapropel formation (type 3; see text for explanation).Variable bottomwater ventilation gives rise to sediment colour variations (a*) where greener hues reflect poorer ventilation and redder hues indicate oxic conditions between sapropels.Top: Eccentricity variations (Laskar, 1990),which reflect orbital control on sapropel deposition (cf. Hilgen, 1991; Rohling and Hilgen, 1991; Rohling, 1994).Modified from Larrasoaña et al.(2003a).

Fig. 24. Photographic illustrations of macroscopic iron sulphide (euhedral pyrite) aggregates that formed in sulphidic microenvironments within otherwise oxic sediments. Brassymillimetric pyrite is evident in the upper middle and right-hand side of (a), and in the upper middle of (b). The brown colour of the background sediments (bioclastic limestone witha micrite ooze matrix) indicates that Fe3+ is widely preserved, while the presence of pyrite indicates that localised microenvironments underwent sulphidic diagenesis. The large piecesof black material in (a) represent recrystallised calcareous macrofossil fragments. Scale: the Australian $2 coin has a 20.5-mm diameter. The images are from polished floor tiles ofunknown provenance.


at lamellar contacts, which gives rise to a strong magnetisation withhigh coercivity. When igneous rocks with such properties are erodedfrom source, detrital ilmenite particles will be deposited in sediments.In most sedimentary environments, these particles would not be

expected to be paleomagnetically significant. However, in relict mag-netic mineral assemblages that lack magnetite, it is reasonable to askwhether such particles can carry a DRM.Wilson and Roberts (1999) de-tected a measurable remanence in single ilmenite crystals, which led


them to suggest that ferrimagnetic microstructural domains mightmake detrital ilmenites responsible for paleomagnetic signals in diage-netically reduced sediments that lack any other detectable remanencecarrying minerals. The lamellar magnetism mechanism of Robinsonet al. (2002) provides a more reasonable explanation for the magneticproperties of detrital ilmenites, but it remains unclear whether smallferrimagnetic particles with lamellar magnetism can give rise to a sedi-mentary DRM. This possibility should be considered and tested in futurestudies of diagenetically reduced sediments.

It is also worth consideringmore broadly the origin of paleomagnet-ic and environmental magnetic signals in diagenetically reduced sedi-ments with relict magnetic mineral assemblages. For example,Hounslow et al. (1995) suggested that dissolution-resistant relict Fe–Cr and Fe–Mn spinels are likely to carry a detrital paleomagnetic signalin sulphidised sedimentary rocks. This is possible, and the reliability ofsuch paleomagnetic signals should be assessed wherever feasible withthe standard tests for paleomagnetic stability (e.g., reversals test, foldtest, etc.), particularly where unusual remanence carriers are likely tobe involved. However, alternative sources of paleomagnetic signals arealso possible in diagenetically reduced sediments. For example, magne-tite particles occur frequently as inclusions within silicate and otherhostminerals (e.g., Feinberg et al., 2005; Tarduno et al., 2006). Such par-ticles are likely to be deposited in sedimentary environments and anymagnetic iron oxide inclusions that they contain will survive sulphidicdiagenesis because the hostminerals are not prone to reductive dissolu-tion. These particles will provide a measurable magnetisation, althoughtheir presence within much larger particles will make them unlikely tobe responsive to an aligning geomagnetic torque. Any such host parti-cles could, therefore, have a paleomagnetically stable net remanence,but it is unlikely to have a geomagnetic origin. Chang et al. (2015) pro-vided direct evidence for preservation of such particles within reduc-tively reduced hemi-pelagic marine sediments offshore of easterncentral Japan. TEM images reveal titanomagnetite nanoparticle inclu-sions, with grain sizes within the tens of nanometre size range, withinlarger silicate host particles (Fig. 26). Ferrimagnetic nanoparticulate in-clusionswithin other host minerals are likely to be important carriers ofmagnetic information from relict magnetic mineral assemblages, butconsiderable care will be needed to interpret such signals.

5. Physical processes and magnetic mineral diagenesis

Discussion of magnetic mineral diagenesis in this paper has mainlyfocused on chemical changes during burial. Physical changes can alsoaffect paleomagnetic recording as well as locking of sedimentarymagnetic fabrics and should strictly be considered as aspects of diage-netic change. Three principal phenomena, associated with sedimentdewatering and compaction, are important: PDRM lock-in, inclinationflattening, and imprinting of sedimentary magnetic fabrics.

5.1. Post-depositional magnetisation lock-in

A DRM is often not recorded in sediments because bioturbation andother processes cause post-depositional remobilisation of sediment par-ticles. A PDRM,which is acquired during earliest burial, is invoked to ex-plain sedimentary magnetisations in such settings (Irving and Major,1964; Kent, 1973; Verosub, 1977). A PDRM is argued to lock-in as a re-sult of progressive consolidation and dewatering, with expulsion of in-terstitial water increasing friction between particles that overcomes

the geomagnetic force that might otherwise impart a realigning torqueon a magnetic particle. Tauxe et al. (2006) argued that the PDRM con-cept is poorly substantiated, although several cases provide strong evi-dence for the existence of a PDRM (Channell and Guyodo, 2004;Sagnotti et al., 2005b; Suganuma et al., 2010, 2011; Snowball et al.,2013). Regardless, while the PDRM concept was proposed 50 yearsago (Irving and Major, 1964), it is fair to say that much more work isneeded to understand this earliest diagenetic process. Key questions re-late to the responses of different sediment types and sediment grainsizes, flocculation and pelletisation of sediment particles, salinity, bio-turbation, compaction, and potential differential contributions toPDRM acquisition from biogenic and detrital magnetic minerals.PDRM acquisition is a significant topic that has been recently reviewedby Roberts et al. (2013b). It is, therefore, not covered in greater detailhere, but it is a subject that requires intensive ongoing research to re-solve substantial questions that limit understanding of how paleomag-netic signals are recorded by sediments.

5.2. Sediment compaction and inclination flattening

Paleomagnetic inclinations in sediments and sedimentary rocks areoften shallower on average than expected for a geocentric axial dipole(GAD) field. This so-called inclination flattening can be explained bytwo factors. First, shallow inclinations have long been reported in asso-ciationwith DRMacquisition if spherical particles roll when they are de-posited onto a sedimentary substrate (e.g., King, 1955; Griffiths et al.,1960; Verosub, 1977; Tauxe, 2005). Second, elongated particles willflatten into the bedding plane of sediments as they are progressivelycompacted during burial (e.g., Anson and Kodama, 1987; Arason andLevi, 1990; Sun and Kodama, 1992). Both phenomena produce inclina-tion flattening of the same form (Anson and Kodama, 1987; Tauxe,2005), which can be described by the “inclination error formula” ofKing (1955):

tan Io ¼ f tan Ie; ð9Þ

where Io and Ie are the observed and expected field inclinations,

Fig. 25. Relict magnetic mineral assemblages that have survived sulphidic dissolution. (a) Euhedral (high titanium) titanomagnetite particle (with X-ray elemental analysis shown belowfor the spot indicated with a red circle), and (b–d) skeletal titanomagnetite particles that consist mainly of residual ilmenite lamellae oriented parallel to the octahedral [111] planes.Magnetite between the ilmenite lamellae has been dissolved. The particles are partially overgrown by euhedral pyrite, as indicated by the X-ray elemental analysis for the yellow circlein (c). (e) Euhedral Cr-magnetite. (f) Partially etched euhedral titanomagnetite particle viewed in a polished section. (g) Partially etched (high titanium) titanomagnetite particle, withX-ray elemental analyses for different spots. All samples except (f) are from magnetic extracts of sediments from Lake Kinneret, Israel (modified from Nowaczyk (2011)), while (f) is fromsediments from Hydrate Ridge, offshore of Oregon (from Larrasoaña et al. (2007)). All images are backscattered electron images obtained with a SEM, except for (e), which is a secondaryelectron image.

Fig. 26. Bright-field TEM image of titanomagnetite nanoparticle intergrowths (indicatedwith arrows)within a silicate hostmineral froma diagenetically reducedmarine sedimentsample from offshore of east-central Japan (Chang et al., 2015). The image was takenfrom the edge of the silicate grain to allow transmission of electrons and imaging of nano-particle inclusions.


respectively, and f is the so-called flattening factor. Flattening associ-ated with deposition is not a diagenetic process, whereascompaction-related flattening is. Anson and Kodama (1987) usedexperimental compaction studies to conclude that compaction-induced inclination flattening is unlikely to occur in natural sedi-ments until burial depths of several hundred metres. In contrast,Arason and Levi (1990) observed noticeable compaction-induced in-clination shallowing at burial depths of 60–85 m. Based on thesestudies, compaction-related inclination shallowing should not be ex-pected at shallow depths above ~60 m, but it is a possible factorbelow this depth. Inclination shallowing is a significant issue in di-verse paleomagnetic applications. For example, systematically shal-low inclinations have led to questioning of the GAD hypothesis andinvocation of persistent ancient non-axial dipole field geometries(e.g., Chauvin et al., 1996; van der Voo and Torsvik, 2001). In tectonicstudies, shallow inclinations have been suggested to indicate majorpaleolatitude differences between the past and present, which im-plies significant poleward tectonic transport (Kent and Irving,2010; Tan et al., 2010; Lippert et al., 2011; Cogné et al., 2013;Muttoni et al., 2013). These issues make it important to havemethods to detect and correct for inclination shallowing.

While it is not possible to distinguish between depositional andcompaction-related inclination flattening, two principal methodsare used to correct for flattening. Jackson et al. (1991) proposed amethod based on measurement of the anisotropy of remanence. Itis effective and has been widely used (e.g., Tan and Kodama, 1998;Weaver et al., 2003; Bilardello and Kodama, 2010), but it is labour-intensive and time-consuming. Tauxe and Kent (2004) proposed analternative method that makes use of the different properties of ex-pected distributions for geomagnetic field directions and flattenedinclinations. Geomagnetic secular variation is expected to give riseto directional distributions that are significantly elongated in theup-down direction at the equator and that are circularly symmetricat the poles. In contrast, inclination flattening gives rise to elongateddistributions depending on the inclination of the applied field andthe value of f. The resulting so-called elongation-inclination (E/I)correction is obtained by inverting measured inclinations using arange of f values to find the E/I pair that is most consistent with a sta-tistical model for secular variation (Tauxe and Kent, 2004; Tauxe,2005). E/I corrections have now been widely used and appear to cor-rect effectively for inclination flattening (e.g., Krijgsman and Tauxe,2004; Tauxe et al., 2008; Muttoni et al., 2013). Detection of flattenedinclinations also obviates any need to invoke non-GAD field configu-rations and poleward tectonic transport (Krijgsman and Tauxe,2004; Tauxe, 2005). Unlike many of the destructive, overprinting orremagnetising effects of diagenesis outlined in this paper, it is possi-ble to correct for compaction effects on paleomagnetic recording.

5.3. Locking of sedimentary magnetic fabrics

Sedimentary magnetic fabrics, which can be detected using theanisotropy of magnetic susceptibility (AMS) or other types of mag-netic anisotropy (Jackson, 1991; Rochette et al., 1992), have beenwidely analysed in tectonic studies (e.g., Borradaile and Tarling,1981; Kissel et al., 1986; Tarling and Hrouda, 1993; Aubourg et al.,1995; Sagnotti et al., 1998; Kanamatsu et al., 2001; Weaver et al.,2004), in analysis of paleocurrent directions (e.g., Rees, 1961, 1965;Hamilton and Rees, 1970; Ellwood and Ledbetter, 1977; Kisselet al., 1997, 1998; Hassold et al., 2006, 2009; Parés et al., 2007),and to detect sedimentary disturbances (e.g., Marino and Ellwood,1978; Rosenbaum et al., 2000; Schwehr and Tauxe, 2003). If elongat-edmagnetic particles are aligned by paleocurrents, themagnetic fab-ric of the sediment will have been preserved from the time ofdeposition. In tectonic studies, it is important to understand thetiming of imprinting of tectonically induced magnetic fabrics. It isgenerally assumed that the local strain field produces a reorientation

of mineral grains shortly after deposition in weakly deformedmudrocks that is indicative of active regional deformation(e.g., Benn, 1994; Borradaile and Tarling, 1981; Parés, 2004; Richteret al., 1993), but this assumption has rarely been tested rigorously.Larrasoaña et al. (2011) demonstrated that the magnetic fabric israpidly locked, within just over 1000 years, in well-dated lake sedi-ments from a compressive tectonic regime in the Kyrgyz Tien Shan.Such rapid post-depositional locking of a magnetic fabric validatesthe use of AMS in tectonic studies and demonstrates theimportance of early diagenetic physical processes in magnetic fabricacquisition.

6. Temperature-dependent diagenetic changes during burial

The diagenetic processes discussed in this paper aremainly appli-cable at temperatures b50 °C. Geothermal gradients in most (non-volcanic) parts of the world are ~25 °C/km (Turcotte and Schubert,2002), which means that the diagenetic processes discussed are gen-erally relevant in sediments that have only been buried to shallowdepths b1–1.5 km. The above considerations are, therefore, appro-priate for most studies of sediment cores. However, burial diagenesiswill be relevant when considering the magnetic properties of sedi-mentary rocks that have been more deeply buried and then ex-humed to the Earth's surface. This situation is likely to be relevantto most outcrops of pre-Tertiary sedimentary rocks, which makesconsideration of the effects of heating during burial widely impor-tant in paleomagnetism (Pullaiah et al., 1975). Magnetic viscosity,where secondary magnetisations are acquired by SD particles,depends on time, temperature and grain size (Néel, 1949). Athermoviscous remanent magnetisation (TVRM) acquired duringburial heating at temperatures up to several hundred degrees Celsiusover prolonged periods of geological time at depths of 5–10 km hasbeen argued to be removable by thermal demagnetisation (Pullaiahet al., 1975). Magnetic overprints due to TVRM acquisition have,therefore, been used for burial temperature paleothermometry(e.g., Middleton and Schmidt, 1982) or to explain remagnetisations(e.g., Kent, 1985). However, as discussed below in Section 7,remagnetisation is generally attributed to chemical processes thatcause mineralogical transformation rather than to TVRM acquisition.

In addition to TVRM acquisition, it is important to consider the pos-sibility of chemical alteration of magnetic minerals during burial. As ob-served during laboratory thermal demagnetisation experiments inpaleomagnetic studies, thermal alteration of non-magnetic mineralsand neoformation of magnetic minerals is common at elevated temper-atures, including formation of magnetite from pyrite (Roberts andPillans, 1993; Passier et al., 2001), siderite (Hirt and Gehring, 1991;Pan et al., 2000), and iron-rich smectites (Hirt et al., 1993). It is, there-fore, reasonable to expect variable stability of magnetic minerals andproduction of new magnetic minerals at the elevated temperatures en-countered by sediments during burial to depths of several km. Fine-grained authigenic magnetite can form at moderate (non-metamor-phic) temperatures during burial diagenesis (Fig. 27) (e.g., Jacksonet al., 1988, 1993; Suk et al., 1990a, 1990b; Banerjee et al., 1997; Katzet al., 1998; Moreau et al., 2005; Aubourg and Pozzi, 2010; Kars et al.,2012, 2014; Blaise et al., 2014). For example, authigenic magnetite canform from pyrite, as evident in pseudomorphing of early diageneticframboidal pyrite (Suk et al., 1990a).

Laboratory heating of thermally immature clays can constrainmagnetic mineral alteration and neoformation under simulatedburial temperatures and pressures, and can inform analysis of sedi-mentary sequences for which independent geothermometric infor-mation is available. This has led to development of a magneticscheme for interpretation of the thermal maturity of sedimentary se-quences (Kars et al., 2014). Abdelmalak et al. (2012) argued thatgoethite will only be present in immature claystones in which burialtemperatures did not exceed 60 °C. Magnetite is argued to form


continuously above 50 °C up to 250 °C (Aubourg and Pozzi, 2010), al-though much of the magnetite produced is ultra-fine-grained andoccurs in the SP state (Kars et al., 2012). This magnetiteneoformation is argued to give rise to burial remagnetisations car-ried by the small fraction of stable SD particles produced during buri-al (Kars et al., 2012), and could also explain the large concentrationsof SP magnetite documented in remagnetised Paleozoic carbonates(e.g., Jackson, 1990; Channell and McCabe, 1994; McCabe andChannell, 1994; Xu et al., 1998; Elmore et al., 2006). Fine-grainedmonoclinic pyrrhotite is argued to form above 150 °C (Aubourg andPozzi, 2010), with magnetite production ceasing above 250 °C andpyrrhotite formation ceasing at ~350 °C (Kars et al., 2014). Thescheme of Kars et al. (2014) is summarised in Fig. 27. As useful asthis model may be for understanding temperature-dependent burialeffects on magnetic minerals, it is non-unique. The presence or ab-sence of goethite, for example, can occur for other reasons, and, as in-dicated in Section 4.2.4.6, monoclinic pyrrhotite can also formdiagenetically at lower temperatures. It is, therefore, necessary totest alternative possibilities for the absence or occurrence of differ-ent magnetic minerals when using the scheme of Kars et al.

(2014); assessment of the processes outlined in this paper should as-sist such efforts.

7. Remagnetisation

“Remagnetisation” is a variable term that is often used to refer to an-cient rocks that have undergone a distinct event that has completelyreset themagnetisation so that the remainingpaleomagnetic signal doc-uments the geomagnetic field at the time of the remagnetisation event.In other cases, remagnetisation can refer to a viscous overprint that doesnot completely reset the magnetisation so that vestiges of the originalmagnetisation can be recovered using techniques such asremagnetisation great circles analysis (e.g., McFadden and McElhinny,1988). In each case, it is generally assumed that the recorded paleomag-netic signal provides information about the geomagnetic field at thetime of the remagnetisation event or at about the time of deposition iftheunderlying signal can be identified below theoverprint.While eithercase is possible, it is important to recognize that remagnetisations canbemuch more subtle than this and that simplistic interpretations can bedangerous. As is the case for paleomagnetic signals carried by haematitein red beds (van der Voo and Torsvik, 2012), remagnetisations carriedby greigite can start from deposition and could stop well within therealm of “early” diagenesis, but could completely disrupt the conven-tional interpretation of paleomagnetic data. It is important to recognisethat remagnetisation is a variable concept. Remagnetisation can occurearly and give rise to paleomagnetic records that are generally adequatefor tectonic applications, but meaningless for analysis of short-periodgeomagnetic field variations. In other cases, remagnetisation canoccur late and traditional paleomagnetic field tests such as the conglom-erate, reversals or fold tests might be the most appropriate means ofidentifying a remagnetisation; it might also be possible to tiethe remagnetisation to a geological event (e.g., folding, fluid flow).Remagnetisations can occur in any environment in which secondarymagnetic minerals form, including oxic (e.g., van der Voo and Torsvik,2012), sulphidic (e.g., Roberts and Weaver, 2005), and methanic envi-ronments (e.g., Weaver et al., 2002; Larrasoaña et al., 2007), or at mod-erate (e.g., Aubourg and Pozzi, 2010; Kars et al., 2012, 2014) or hightemperatures (e.g., Appel et al., 2012; Abrajevitch et al., 2014).

Remagnetisations have been widely reported, particularly in Paleo-zoic and Mesozoic carbonates (see reviews by McCabe and Elmore(1989); Jackson and Swanson-Hysell (2012) and van der Voo andTorsvik (2012)). Remagnetised carbonates often contain high concen-trations of authigenic SP magnetite, which could be produced by burialdiagenesis of the type proposed by Aubourg and Pozzi (2010) and Karset al. (2012, 2014). Although SP particles do not contribute to the natu-ral remanent magnetisation, their presence in large concentrations inremagnetised carbonates, together with stable SD particles, producescharacteristic wasp-waisted hysteresis behaviour (Jackson, 1990;Roberts et al., 1995; Tauxe et al., 1996) that contrasts with that ofnon-remagnetised carbonates (e.g., Roberts et al., 2013a). Such hystere-sis signatures have been argued to provide a fingerprint for carbonateremagnetisation (Jackson, 1990; McCabe and Channell, 1994; Channelland McCabe, 1994; Jackson and Swanson-Hysell, 2012; van der Vooand Torsvik, 2012). As is usually the case, various processes can producenon-unique hysteresis properties, which complicates such simple testsand makes it important to apply a range of tests for remagnetisations.Multiple mechanisms have been proposed for carbonateremagnetisations, including continent-scale orogenic fluid flow(Oliver, 1986), TVRM acquisition (Kent, 1985), burial to temperaturesthat cause thermal alteration of pyrite to magnetite (e.g., Suk et al.,1990a, 1990b; Banerjee et al., 1997), transformation of iron-rich smec-tite to iron-free illite + magnetite (McCabe et al., 1989; Katz et al.,1998; Tohver et al., 2008), and pressure solution (Zegers et al., 2003).Most of thesemechanisms occur above the ~50 °C threshold consideredhere for diagenesis in thermally immature sediments.

Fig. 27. Temperature-dependent magnetic mineral alteration and formation with burial.Burial temperature calibrations are from vitrinite reflectance (Ro) data. The oil and gaswindows are indicated. Modified from Kars et al. (2014).


8. Outstanding questions concerning magnetic mineral diagenesis

Many issues concerning magnetic mineral diagenesis are well un-derstood, while many others remain under-constrained. The chemicalreactions involved in diagenetic magnetic mineral modification arewell understood throughout the range of environments from oxic tosulphidic, although processes in the methanic zone are less well under-stood. Reaching the level of knowledge that exists for shallower diage-netic zones will require detailed multi-disciplinary work on processesin the methanic zone that involve (bio-)geochemistry, microbiology,sedimentology, reactive transport modelling and mineral magnetism.Much also remains to be done to understand the role of microbes inthe myriad reactions that occur in the full range of diagenetic environ-ments. Microbes are crucially important in determining the rate atwhich many magnetic mineral formation, transformation and destruc-tion processes proceed, yet fundamental questions remain about theirpresence, their interactions with broader microbial consortia, and theextent to which they catalyse or control important processes. Advancesin geomicrobiology are likely to play an increasingly important role inproviding a better understanding of diagenetic processes. In particular,it is unknown to what extent extracellular magnetic by-products of mi-crobial metabolism, which involve formation of both magnetite andgreigite, aremagnetically important for sediments in nitrogenous, ferru-ginous and sulphidic environments. Future work is needed to establishquantitatively the importance of these processes, including the degreeof preservation of these magnetic particles in the geological record.

Despite the fact that sedimentary pyrite formation has been studiedfor decades, the chemical pathways by which iron sulphides form, in-cluding the importance of ferrimagnetic greigite, remain controversial(e.g., Schoonen, 2004). Answers to these questions are likely to comefrom detailed geochemical andmicrobiological work, and are necessaryto better understand theenvironmental andpaleomagnetic implicationsof greigite formation and preservation in the geological record. Recogni-tion of the widespread importance of greigite in the geological recordhas largely been driven by the sensitivity of magnetic measurementsfor detecting small quantities of magnetic minerals (e.g., Roberts et al.,2011b), so there is much that magnetic analysis can contribute toassessing these questions.

Recent work has established the importance of magnetitemagnetofossils in the geological record (Roberts et al., 2012). This ischanging a situation whereby, until recently, magnetite magnetofossilswere relatively rarely reported in the pre-Quaternary geological record(Kopp and Kirschvink, 2008). Key questions that remain to be wellconstrained concern the depth at which magnetotactic bacteria lived(i.e., within the water column or sediment). If the bacteria lived withinsediment, particularly below the surface mixed layer, it will be neces-sary to understand whether the associated biogeochemical remanentmagnetisation produced by such biogenic particles (Tarduno andWilkison, 1996; Tarduno et al., 1998) is widely responsible for paleo-magnetic signals in the geological record (Roberts et al., 2013a;Larrasoaña et al., 2014). It is also expected that, in the coming years,greigite magnetofossils will prove to be much more commonly pre-served in sulphidic sediments (Chang et al., 2014).

Recent work has also improved our understanding of the effects ofburial diagenesis on magnetic minerals (e.g., Aubourg and Pozzi, 2010;Kars et al., 2014). Ongoingwork is needed to test and refine the schemeof Kars et al. (2014) to assess the limits of its applicability. If it proves tocorrectly represent the effects of burial on magnetic minerals, it shouldbecomewidely important inmagnetic analyses of moderately to deeplyburied sediments and in geothermometric applications.

Relictmagnetic mineral assemblages (ilmenite, Cr spinels, etc.) havebeen widely reported in sulphidic diagenetic environments in whichdetrital magnetic minerals have largely been removed by reductive dis-solution (e.g., Roberts and Turner, 1993; Hounslow et al., 1995;Hounslow, 1996; Wilson and Roberts, 1999; Nowaczyk, 2011). Thecommonpresence of ilmenite inmagnetic extracts from such sediments

raises the question as towhether the lamellarmagnetismmechanismofRobinson et al. (2002) can give rise to paleomagnetic signal recordingby fine-grained ilmenite particles in diagenetically reduced sediments.This possibility is yet to be rigorously tested.

Finally, several questions remain open in relation to the effects of sil-ica diagenesis on magnetic minerals. Florindo et al. (2003) proposedthat silica diagenesiswill give rise tomagnetite dissolution. This propos-al has not been widely tested and conflicting results have been pub-lished. Even if this relationship is not valid, recent work suggests thatdiagenetic iron oxide precipitation can be important in chert formation(Meister et al., 2014), which provides an important avenue for futureinvestigations.

9. Conclusions

Magnetic mineral diagenesis plays a key role in the global iron cycle.The sensitivity of mineral magnetic properties to important processesmeans that magnetic analyses can be useful for understanding theglobal iron cycle if the processes at play can be well constrained andunderstood. Understanding the extent of diagenetic alteration and/orauthigenicmagneticmineral formation is also fundamentally importantfor interpretation of environmental magnetic and paleomagnetic sig-nals. These effects can range from subtle to pervasive; the extent oftheir influence needs to be understood in virtually all magnetic analysesof sediments and sedimentary rocks.

Many possibilities exist for using magnetic mineral diagenesis, par-ticularly non-steady state variations that reflect climate variability, toelucidate important environmental processes (e.g., Larrasoaña et al.,2003a; Blanchet et al., 2009; Nowaczyk et al., 2012; Chang et al.,2015), although the possibility of diagenesis obscuring importantclimate processes should always be considered (e.g., Abrajevitch et al.,2009). The range of diagenetic processes that affect magnetic mineralsis extensive and has not been systematically summarised before in asingle source despite the large literature on various aspects of magneticmineral diagenesis. It is hoped that the relatively detailed treatment ofmagnetic mineral diagenesis provided here will assist researchers inunravelling the effects of magnetic mineral diagenesis in a range of set-tings, and that highlighting significant under-constrained issues willlead to concerted efforts to better understand key processes.

Acknowledgements

This work benefitted from the support of the Australian ResearchCouncil (through grant DP120103952). I thank colleagues who gener-ously provided images that are reproduced in this paper, includingAlexandra Abrajevitch, Liao Chang, Ramon Egli, Chorng-Shern Horng,Wei-Teh Jiang, Juan-Cruz Larrasoaña, Kais Mohamed, NorbertNowaczyk, Alan Palmer, Yongxin Pan, Brad Pillans, Chris Rowan, JimWatson, and RichardWeaver. Xiang Zhao provided assistance with pro-ducing some of the figures. I have benefitted over the years fromdiscus-sions with many colleagues about magnetic mineral diagenesis,particularly Bob Karlin, Rich Reynolds, Steve Lund, Chorng-ShernHorng, Shuh-Ji Kao, Mark Dekkers, Juan-Cruz Larrasoaña, Chris Rowan,Liao Chang, Leonardo Sagnotti, Fabio Florindo, Rob Raiswell, LianeBenning, Martin Palmer, and Ken Nealson. The paper also benefittedconsiderably from the helpful and insightful reviews of Sabine Kasten,Mark Dekkers, and Toshi Yamazaki. I am indebted to these colleagues,but all errors and misconceptions remain my own.

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