# Reviews and syntheses: Review of proxies for low-oxygen paleoceanographic reconstructions

Babette A.A. Hoogakker, Catherine Davis, Yi Wang, Stephanie Kusch, Katrina Nilsson-Kerr, Dalton Hardisty, Allison Jacobel, Dharma Reyes Macaya, Nicolaas Glock, Sha Ni, et al.

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Babette A.A. Hoogakker, Catherine Davis, Yi Wang, Stephanie Kusch, Katrina Nilsson-Kerr, et al. Reviews and syntheses: Review of proxies for low-oxygen paleoceanographic reconstructions. Biogeosciences, 2025, 22 (4), pp.863-957. 10.5194/bg-22-863-2025. hal-05073795

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# Reviews and syntheses: Review of proxies for low-oxygen paleoceanographic reconstructions

Babette A.A. Hoogakker<fnref n="1" />, Catherine Davis<fnref n="2" />, Yi Wang<fnref n="3" />, Stephanie Kusch<fnref n="4" />, Katrina Nilsson-Kerr<fnref n="5" />,
Dalton S. Hardisty<fnref n="6" />, Allison Jacobel<fnref n="7" />, Dharma Reyes Macaya<fnref n="1" />,<fnref n="8" />,<fnref n="9" />, Nicolaas Glock<fnref n="10" />, Sha Ni<fnref n="10" />, Julio Sepúlveda<fnref n="11" />,
Abby Ren<fnref n="12" />, Alexandra Auderset<fnref n="13" />, Anya V. Hess<fnref n="14" />, Katrin J. Meissner<fnref n="15" />, Jorge Cardich<fnref n="16" />, Robert Anderson<fnref n="17" />,
Christine Barras<fnref n="18" />, Chandranath Basak<fnref n="19" />, Harold J. Bradbury<fnref n="20" />, Inda Brinkmann<fnref n="21" />, Alexis Castillo<fnref n="9" />,
Madelyn Cook<fnref n="22" />,<fnref n="23" />, Kassandra Costa<fnref n="24" />, Constance Choquel<fnref n="2" />,<fnref n="1" />, Paula Diz<fnref n="25" />, Jonas Donnenfield<fnref n="26" />, Felix J. Elling²⁷,
Zeynep Erdem²⁸, Helena L. Filipsson<fnref n="21" />, Sebastián Garrido<fnref n="1" />, Julia Gottschalk²⁹, Anjaly Govindankutty Menon<fnref n="10" />,
Jeroen Groeneveld³⁰, Christian Hallmann³¹,³², Ingrid Hendy<fnref n="22" />, Rick Hennekam³³, Wanyi Lu³⁴,
Jean Lynch-Stieglitz³⁵, Lélia Matos³⁶,³⁷, Alfredo Martínez-García³⁸, Giulia Molina³⁶,³⁷, Práxedes Muñoz³⁹,
Simone Moretti³⁸, Jennifer Morford⁴⁰, Sophie Nuber<fnref n="12" />, Svetlana Radionovskaya⁴¹, Morgan Reed Raven⁴²,
Christopher J. Somes⁴³, Anja S. Studer⁴⁴, Kazuyo Tachikawa⁴⁵, Raúl Tapia³⁰, Martin Tetard⁴⁶, Tyler Vollmer³⁵,
Xingchen Wang⁴⁷, Shuzhuang Wu⁴⁸, Yan Zhang⁴⁹, Xin-Yuan Zheng⁵⁰, and Yuxin Zhou⁴²

<footnote n="1">¹The Lyell Centre, Heriot-Watt University, Edinburgh, EH14 4AP, UK</footnote>

<footnote n="2">²Department of Marine, Earth and Atmospheric Sciences, North Carolina State University, Raleigh, NC 27695, USA</footnote>

<footnote n="3">³Department of Earth and Environmental Sciences, Tulane University, New Orleans, LA 70118, USA</footnote>

<footnote n="4">⁴Institute of Marine Sciences, University of Quebec Rimouski, Rimouski, QC G5L 3A1, Canada</footnote>

<footnote n="5">⁵Department of Earth Sciences, University of Bergen and Bjerknes Centre for Climate Research, Bergen, 5007, Norway</footnote>

<footnote n="6">⁶Department of Earth and Environmental Sciences, Michigan State University, East Lansing, MI 48824, USA</footnote>

<footnote n="7">⁷Department of Earth and Climate Sciences, Middlebury College, Middlebury, VT 05753, USA</footnote>

<footnote n="8">⁸Center for Marine Environmental Sciences, University of Bremen, 28359 Bremen, Germany</footnote>

<footnote n="9">⁹Centro de Investigación de Estudios Avanzados del Maule, Universidad Católica del Maule, Campus San Miguel, Talca, 3460000, Chile</footnote>

<footnote n="10">¹⁰Institute for Geology, Department of Earth System Sciences, University of Hamburg, 20146 Hamburg, Germany</footnote>

<footnote n="11">¹¹Department of Geological Sciences and Institute of Arctic and Alpine Research, University of Colorado Boulder, Boulder, CO 80309, USA</footnote>

<footnote n="12">¹²Department of Geosciences, National Taiwan University, Taipei 106, Taiwan</footnote>

<footnote n="13">¹³School of Ocean and Earth Science, University of Southampton, Southampton, SO14 3ZH, UK</footnote>

<footnote n="14">¹⁴Department of Earth and Planetary Sciences, Rutgers, the State University of New Jersey, New Brunswick, NJ 08854, USA</footnote>

<footnote n="15">¹⁵Climate Change Research Centre and ARC Centre of Excellence for Climate Extremes, University of New South Wales, Sydney, NSW 2052, Australia</footnote>

<footnote n="16">¹⁶CIDIS - Facultad de Ciencias e Ingeniería, Universidad Peruana Cayetano Heredia, Lima 15102, Peru</footnote>

<footnote n="17">¹⁷Lamont-Doherty Earth Observatory of Columbia University, Palisades, NY 10964, USA</footnote>

<footnote n="18">¹⁸Univ Angers, Nantes Université, Le Mans Univ, CNRS, UMR 6112, Laboratoire de Planétologie et Géosciences, 49000 Angers, France</footnote>

<footnote n="19">¹⁹Department of Earth Sciences, University of Delaware, Newark, DE 19716, USA</footnote>

<footnote n="20">²⁰Department of Earth, Ocean and Atmospheric Sciences, University of British Columbia, Vancouver, V6T 1Z4, Canada</footnote>

<footnote n="21">²¹Department of Geology, Lund University, Lund, 223 63, Sweden</footnote>

<footnote n="22">²²Department of Earth and Environmental Sciences, University of Michigan, Ann Arbor, MI 48109, USA</footnote>

<footnote n="23">²³Department of Geosciences, University of Arizona, Tucson, AZ 85721, USA</footnote>

<footnote n="24">²⁴Department of Geology and Geophysics, Woods Hole Oceanographic Institution, Falmouth, MA 02543, USA</footnote>

<footnote n="25">²⁵Centro de Investigación Mariña, XM1, Universidade de Vigo, Vigo, 36310, Spain</footnote>

<footnote n="26">²⁶College of Earth, Ocean, and Atmospheric Sciences, Oregon State University, Corvallis, OR 97331, USA</footnote>
<sup>27</sup>Leibniz-Laboratory for Radiometric Dating and Stable Isotope Research, Christian-Albrecht University of Kiel, 24118 Kiel, Germany

<sup>28</sup>Department of Marine Microbiology and Biogeochemistry, NIOZ Royal Netherlands Institute for Sea Research, 't Horntje (Texel), Den Hoorn, 1797 SZ, the Netherlands

<sup>29</sup>Institute of Geosciences, University of Kiel, 24118 Kiel, Germany

<sup>30</sup>Institute of Oceanography, National Taiwan University, Taipei 106, Taiwan

<sup>31</sup>GFZ German Research Centre for Geosciences, 14473 Potsdam, Germany

<sup>32</sup>Institute of Geosciences, University of Potsdam, 14476 Potsdam, Germany

<sup>33</sup>Department of Ocean Systems, NIOZ Royal Netherlands Institute for Sea Research, Den Burg (Texel), 1790 AB, the Netherlands

<sup>34</sup>State Key Laboratory of Marine Geology, Tongji University, Shanghai 200092, China

<sup>35</sup>School of Earth and Atmospheric Sciences, Georgia Institute of Technology, Atlanta, GA 30332, USA

<sup>36</sup>Centre of Marine Sciences (CCMAR), University of Algarve, Faro, 8005-139, Portugal

<sup>37</sup>Marine Geology and Georesources Division, Portuguese Institute for Sea and Atmosphere (IPMA), Lisbon, 1495-165, Portugal

<sup>38</sup>Max Planck Institute for Chemistry, 55128 Mainz, Germany

<sup>39</sup>Departamento de Biología Marina, Universidad Católica del Norte, Coquimbo, 1780000, Chile

<sup>40</sup>Chemistry Department, Franklin and Marshall College, Lancaster, PA 17604, USA

<sup>41</sup>Department of Earth Sciences, University of Cambridge, Cambridge, CB23EQ, UK

<sup>42</sup>Department of Earth Science, University of California, Santa Barbara, CA 93106, USA

<sup>43</sup>GEOMAR Helmholtz Centre for Ocean Research Kiel, 24148 Kiel, Germany

<sup>44</sup>Department of Environmental Sciences, University of Basel, 4056 Basel, Switzerland

<sup>45</sup>Aix Marseille Univ, CNRS, IRD, INRAE, Coll France, CEREGE, Aix-en-Provence, 13545, France

<sup>46</sup>GNS Science, Lower Hutt, 5011, Aotearoa/New Zealand

<sup>47</sup>Department of Earth and Environmental Sciences, Boston College, 1 Fairway Drive, Boston, MA 02467, USA

<sup>48</sup>Institute of Earth Sciences, University of Lausanne, 1015 Lausanne, Switzerland

<sup>49</sup>Ocean Sciences Department, University of California, Santa Cruz, CA 95064, USA

<sup>50</sup>Department of Earth and Environmental Sciences, University of Minnesota, Minneapolis, MN 55455, USA

**Correspondence:** Babette A.A. Hoogakker (b.hoogakker@hw.ac.uk) and Catherine Davis (cdavis24@ncsu.edu)

Received: 11 December 2023 – Discussion started: 9 January 2024

Revised: 25 October 2024 – Accepted: 14 November 2024 – Published: 18 February 2025

**Abstract.** A growing body of observations has revealed rapid changes in both the total inventory and the distribution of marine oxygen over the latter half of the 20th century, leading to increased interest in extending oxygenation records into the past. The use of paleo-oxygen proxies has the potential to extend the spatial and temporal range of current records, constrain pre-anthropogenic baselines, provide datasets necessary to test climate models under different boundary conditions, and ultimately understand how ocean oxygenation responds beyond decadal-scale changes. This review seeks to summarize the current state of knowledge about proxies for reconstructing Cenozoic marine oxygen: sedimentary features, sedimentary redox-sensitive trace elements and isotopes, biomarkers, nitrogen isotopes, foraminiferal trace elements, foraminiferal assemblages, foraminiferal morphometrics, and benthic foraminiferal carbon isotope gradients. Taking stock of each proxy reveals some common limitations as the majority of proxies functions best at low-oxygen concentrations, and many reflect multiple environmental drivers.

We also highlight recent breakthroughs in geochemistry and proxy approaches to constraining pelagic (in addition to benthic) oxygenation that are rapidly advancing the field. In light of both the emergence of new proxies and the persistent multiple driver problem, the need for multi-proxy approaches and data storage and sharing that adhere to the principles of findability, accessibility, interoperability, and reusability (FAIR) is emphasized. Continued refinements of proxy approaches and both proxy–proxy and proxy–model comparisons are likely to support the growing needs of both oceanographers and paleoceanographers interested in paleo-oxygenation records.

# 1 Introduction

Dissolved oxygen in the ocean is necessary to sustain aerobic life and control biogeochemical processes and is closely
linked to carbon remineralization, export, and storage. Oxygen in the ocean has declined since at least the mid-20th century. This decrease has been observed in estuaries and coastal regions (Diaz and Rosenberg, 2008; Rabalais, 2009; Rabalais et al., 2010; Conley et al., 2011), continental shelves, and the open ocean (Schmidtko et al., 2017; Chan et al., 2008; Bograd et al., 2008; Breitburg et al., 2018; Keeling et al., 2010; Levin, 2018; Stramma et al., 2008; Stramma et al., 2010). Direct measurements of oxygen have only been routine for a few decades at most but, even so, are spatially limited. Inaccessible subsurface regions and open ocean features, such as oxygen minimum zones (OMZs), are especially difficult to monitor. Thus, proxies are required to extend modern records and investigate long-term drivers of deoxygenation.

Drivers of ocean deoxygenation include (1) ocean warming, causing decreasing oxygen solubility in seawater and increasing remineralization rates; (2) increased productivity, leading to higher subsurface oxygen utilization during respiration; and (3) decreased ventilation due to changes in circulation or stratification (Keeling et al., 2010; Breitburg et al., 2018). These drivers can influence ocean deoxygenation on different timescales and to different degrees. Warming is a key driver of modern deoxygenation in the open ocean and in coastal systems (Schmidtko et al., 2017; Levin, 2018; Rabalais et al., 2010). In coastal systems, anthropogenic nutrient increases (eutrophication) from activities such as sewage efflux and fertilizer input are frequently the primary cause of deoxygenation on short timescales (Rabalais et al., 2010; Breitburg et al., 2018). Productivity changes can also be important in driving decadal- (Deutsch et al., 2011, 2014) to centennial- and longer-scale changes in open ocean settings (e.g. Hendy et al., 2004). Ventilation changes may act across different scales of space and time. For example, deoxygenation induced by stratification can be variable on timescales of days to years and beyond, especially in coastal regions and restricted basins (reviewed in Rabalais et al., 2010). However, seawater oxygen content is also responsive to ventilation changes on centennial, millennial, and longer timescales, associated with changes in deep-water sources, upwelling, overturning circulation, ocean gateway dynamics, and the geometry of whole ocean basins (Hoogakker et al., 2015; Fyke et al., 2015; Cardich et al., 2019; Auderset et al., 2022; Hess et al., 2023; Khon et al., 2023).

Climate models indicate that a decrease in dissolved oxygen concentrations will continue for hundreds to thousands of years into the future (Bahl et al., 2019; Kwiatkowski et al., 2020; Oschlies, 2021; Gulev et al., 2021). The combined effect of future warming and seawater oxygen depletion could have adverse impacts on the marine environment, potentially culminating in a mass extinction rivalling that in Earth's past (Penn and Deutsch, 2022). The latest state-of-the-art coupled climate models capture the global observational trend in the upper ocean within the conservative end of uncertainty levels (Takano et al., 2023), which are high due to spatiotemporal data sparsity (Ito et al., 2017). However, models still underes-

timate deoxygenation in the deep ocean and do not reproduce the observed patterns in the tropical thermocline (Oschlies et al., 2018; Kwiatkoski et al., 2020), where persistent oxygen-deficient zones exist. This mismatch is likely due to unresolved circulation, mixing, and transport processes; misrepresentation of respiratory oxygen demand; missing geochemical feedback mechanisms; and an insufficient simulation length to reach equilibrium in the deep ocean (Oschlies et al., 2018). To better constrain biological and physical processes in the ocean and improve their representation in models (see Supplement for details), we need dedicated observational programmes. We also need proxy-based oxygen reconstructions from the geologic past, when the climate system was different from that of the present day, to test numerical models and to improve process understanding.

Interest in seawater oxygen proxies is increasing, partly due to current trends of ocean deoxygenation and uncertainties about the future at different timescales. A methodological overview of proxies was included in Moffitt et al. (2015). Since this review was published, methodological developments, updates, and insights have emerged that were not captured previously or were applied to older sediments. The present review is limited to proxies that can be applied throughout the Cenozoic (i.e. the last 66 million years), although we briefly touch upon some well-studied earlier examples, such as Cretaceous oceanic anoxic events (OAEs). The focus on the Cenozoic, when our oceans were overall well oxygenated, allows for an investigation of scenarios and timescales most immediately relevant to informing the future.

Extending modern records into the past provides baselines for pre-industrial marine oxygen content and the necessary data to test climate models under boundary conditions that are different from today's and improve process understanding. While the past is only a partial analogue to the future, it can provide a portfolio of oxygen scenarios to constrain future projections. This is especially the case for past climate episodes that were characterized by greenhouse gas concentrations similar to projected future levels. In step with the growing interest in modern and future ocean oxygen, there has been rapid proxy development over the past few decades. Implementation of new technologies and a burgeoning interest in paleo-oxygenation have led to an influx of new proxies and refocusing and advancement of established proxies.

This review aims to provide an overview of the current state of proxy development at a pivotal moment for the field. We summarize the major classes of proxies for the benefit of both new and experienced paleoceanographers and those working in adjacent fields. It is our hope that an introduction to and update of the suite of available proxies will increase their utility for those interested in marine oxygen research. Moreover, we hope that a clear discussion of current limitations and future directions can pave the way for improving the tools at our disposal for generating new paleo-oxygenation records.
## 2 Proxies

Proxies provide indirect representations of environmental variables in circumstances where they cannot be measured directly, such as the geological past. Examples include seawater temperature, pH, and dissolved oxygen. A proxy is a measurable physical or chemical variable that is conserved in a natural climate archive and allows us to infer information about the variable of interest in a qualitative or quantitative manner. To build a useful proxy, it is important to understand how the proxy relates to the variable of interest and what other environmental parameters might influence the proxy pre- and post-deposition in sediments. This involves understanding the biology (especially if the proxy is captured in fossil and organic material), chemistry, and physics of both proxy and sedimentary systems.

Paleo-oxygen proxies are generally developed and calibrated through a combination of theoretical, empirical, and experimental approaches. Examples of theoretical approaches to proxy development can be subfields of geochemistry, such as inorganic and organic geochemistry of sediments and biogenic calcites. For example, a theoretical understanding of redox potential can lead to robust predictions about concentrations of elements and ions across oxygen gradients and thus whether one would expect redox-sensitive elements to be found in higher or lower abundance in sediments or biogenic minerals. Theoretical approaches generally require empirical validation as many complexities remain difficult to quantify and/or model. For example, redox-associated chemistry and incorporation of products into biogenic minerals are biologically mediated and influenced by other environmental (e.g. temperature) variables and tax-specific dynamics related to their life cycle, metabolism, and ontogeny. As a result, theoretical approaches are usually limited to the identification of proxies of interest and qualitative predictions.

The use of recently deposited sediments on the seafloor (frequently referred to as “core tops”) recovered across natural oxygen gradients is the most frequent empirical approach. Core-top calibrations can be critical for proxies that require timescales or depositional environments that are difficult to replicate in a laboratory setting, such as foraminiferal assemblages, sedimentary features, and sedimentary trace metals. This approach has the benefit of testing how a proxy manifests in the complex natural environment. One key limitation is the need to deconvolve highly correlated environmental controls, such as productivity, organic carbon content, and oxygen, which are classically difficult to disentangle as drivers of foraminiferal assemblages (Gooday, 2003). This may also impact most other proxies, including the isotopic composition of nitrogen ($\delta^{15}$N) and organic matter. The second key limitation of core-top calibrations is the no-analogue problem; extrapolation beyond modern examples may be required to describe paleo-oxygen environments which are unlike current conditions, particularly during the more extreme

events of ocean deoxygenation found in the geologic record. Furthermore, core tops are not always modern in age due to extremely low sedimentation rates, dissolution, and hibata caused by winnowing currents or active tectonic activity (Mekik and Anderson, 2018; Erdem et al., 2016). Sediment trap studies and plankton tows are other examples of important yet less frequently used empirical approaches.

Experimental approaches are often considered the “gold standard” for quantitative calibration of single-driver proxies. As of now, most paleo-oxygen proxies are qualitative or semi-quantitative. Experimental approaches have the benefit of allowing for single controlling variables to be isolated and have so far been used to the greatest effect in biogenic calcites. However, there are a few drawbacks to this approach. The first is that proxies are removed from the complexities of the natural environment; thus results must be validated with field observations where possible. In other cases, the timescales (e.g. sedimentary features, sedimentary trace metals) or complex initial conditions (e.g. biomarker assemblages) necessary to replicate natural observations are difficult or impossible to generate in a laboratory setting.

## Proxy material

Sediments provide the backbone for any marine paleoenvironmental reconstruction along with its preserved or fossilized biogenic materials. This can include morphologically identifiable skeletal material, such as foraminiferal tests, diatom frustules, or “molecular fossils”.

Our review of the various proxy methods is split into a traditional overview of sediments as proxy carriers (Sect. 4), followed by a discussion of sedimentary redox trace elements and isotopes (Sect. 5), organic proxies (Sect. 6), and nitrogen isotopes (Sect. 7). Following this, we discuss foraminiferal trace elements (Sect. 8), foraminiferal assemblages (Sect. 9), foraminiferal morphometrics (Sect. 10), and benthic foraminiferal carbon isotope gradients (Sect. 11) that use foraminifera as proxy carriers, which are introduced below.

Foraminifera (kingdom – Chromista, infrakingdom – Rhizaria, order – Foraminiferida) are amoeboid protists characterized by a cytoplasmic body and a shell or “test” comprising one or more interconnected chambers. The test wall can be made of agglutinated particles, organic material, or biomineralized crystals of calcite, aragonite, or (rarely) silica (Loeblich and Tappan, 1988). Calcareous tests, in particular, are frequently preserved in marine sediments after death or reproduction (Debenay, 2012). As a result, a rich fossil record of calcareous foraminifera extends from the Cambrian into the present (Loeblich and Tappan, 1988; Sen Gupta, 2003; Debenay, 2012). Foraminifera have colonized a diversity of environments. The majority are benthic, where they occupy virtually every water depth and substrate, on and into the sediment (e.g. Vickerman, 1992; Gooday, 2003; Sen Gupta, 2003). Others are planktic, with habitats ranging
from the ocean's surface into the mesopelagic (Schiebel and Hemleben, 2017). As a result, foraminifera can offer a near-continuous record of ecological succession, with individual shells capturing environmental conditions over their weeks-to years-long lifespans.

The most diverse and abundant living group of foraminifera are the Rotaliida. The calcification and morphology of this group are different from other groups (e.g. Miliolida, Nodosariida, and Robertinida) (De Nooijer et al., 2023). Calcareous hyaline foraminifera, Rotaliida, diversified during the Cretaceous and are the basis of several proxy methods using foraminifera (Sects. 8–11), limiting these to the late Cretaceous until recent (Loeblich and Tappan, 1988; Kaiho, 1994). Other groups, such as Nodosariidae, are sometimes used in paleoenvironmental reconstructions but have differing calcification mechanisms, resulting in marked differences in geochemistry from the more common Rotaliida (De Nooijer et al., 2023; Pacho et al., 2023). Fully planktic Rotaliida foraminifera evolved from benthic orders starting in the Middle Jurassic (BouDagher-Fadel, 2015), meaning that their proxy applications are more limited in time than those of benthic foraminifera.

Living benthic foraminifera are found both on the sediment surface (epifaunal) and throughout at least the top 10 cm of the sediment (infaunal), although the proportion of the total living population decreases rapidly with increasing depth (Corliss, 1985, 1991). Moreover, some benthic foraminifera can migrate vertically within sediments, with their habitat depths and position in the sediment influenced by the organic matter flux and availability of resources such as oxygen (Bernhard, 1992; Barmawidjaja et al., 1995; Loubere et al., 1993; Linke and Lutze, 1993; Jorissen et al., 1995; Geslin et al., 2004). Planktic foraminifera inhabit the water column above the seafloor, with most species having a near-surface habitat (Schiebel and Hemleben, 2017). Benthic foraminifera can therefore record variations in bottom water oxygen concentrations in locations where the seafloor is in a low-oxygen environment. Planktic foraminifera proxies can provide information about both open ocean oxygen conditions and the extent of low-oxygen zones, such as OMZs.

### 3 Terminology and units

As is evident from the later discussion of the different proxies, the nomenclature used to define different oxygenation zonations has historically been inconsistent and confusing (Canfield and Thamdrup, 2009). This can be ascribed in part to the interdisciplinarity of modern oxygen research. Classically, geochemists define an oxic zone as one supporting aerobic metabolism, followed by an oxygen-depleted zone, sometimes referred to as suboxic, where metabolism is supported by nitrate, manganese (Mn), and iron (Fe) reduction, and an anoxic zone, where metabolism is supported by sulfate reduction and methanogenesis (Froelich

et al., 1979; Berner, 1981). However, this scheme has been regarded as confusing and contradictory by Canfield and Thamdrup (2009), who proposed instead using terminal electron acceptors and respiration processes to define chemical/metabolic zones (Fig. 1). Ecologists and biologists have frequently focused on oxygen levels associated with negative outcomes for aerobic organisms (fish, crustaceans, etc.) and have defined a sublethal threshold and lethal oxygen concentrations, which vary greatly among taxa and may be influenced by other factors such as temperature (Vaquer-Suyer and Duarte, 2008). This sublethal threshold is referred to as hypoxia. It leads to mortality events, losses of biodiversity, habitat reduction, predation potential, and disruption of life cycles (Service, 2004; Rabalais et al., 2002). Some literature additionally uses the term suboxia as an intermediate between either oxic and hypoxic or hypoxic and anoxic. The dearth of observational oxygen data at the full range of spatial and temporal scales applicable to either geochemical or ecological systems further complicates definitions in terminology. We further note the use of oxygen-deficient zone (ODZ), a term which is primarily used to describe a region where oxygen is low enough to allow for denitrification or other anaerobic metabolism. The term OMZ is used more broadly to refer to regions of notably low oxygen at a variety of thresholds, frequently defined by dissolved oxygen content. With this usage, all ODZs are also OMZs; however the reverse is not always the case, and both terms are used here in different contexts. To avoid confusion between the different terms used, an illustrative figure is provided to give a sense of the zonation, chemical speciation, and metabolic processes, alongside the “oxygen working” range of the different proxies (Fig. 1). While anoxic is consistently used to describe no (or undetectable) oxygen, other terms are used by different authors to describe different oxygen ranges. This is represented by varying opacity in Fig. 1 to represent the oxygen ranges often associated with these terms. Similarly, units for dissolved oxygen vary markedly, most frequently reported as mL L⁻¹, mg L⁻¹, percent saturation, or µmol kg⁻¹. Unfortunately, conversions are not always straightforward as seawater density needs to be considered. Here, we favour µmol kg⁻¹ for consistency but occasionally reference oxygen in other units when referring to previously published work.

### 4 Sediment properties as proxies

Reconstructions of past marine environments rely on sediment samples from deep sea cores or outcrops of uplifted marine sediments. Sedimentary observations form the backbone of metadata essential to support the growing arsenal of proxies employed to define Earth's biogeochemical evolution. In particular, quantitative mineralogy and lithologic descriptions should accompany sample archives to support existing and future geochemical proxy measurements and interpretations. This is critical as a given proxy may only
<figure><img src="image_3.png" /><figcaption>**Figure 1.** Overview of oxygen “stage” nomenclature used in this review. Panel (a) shows the ranges most often associated with the descriptive terms OMZ, oxygen-deficient zone (ODZ), anoxia, dysoxia/hypoxia, and anoxia in seawater. In (b), oxygen concentrations are shown on a log-linear scale along with a simplified schematic of several proxy-relevant components of other redox-sensitive reactions. Chemical concentrations other than oxygen are non-dimensional, but all relate to scales in both (a) and (c). The redox ladder is modified from Canfield and Thamdrup (2009). Panel (c) shows the ranges of oxygen and/or redox chemistries over which different proxy types can be used to reconstruct paleoenvironments, based on proxies applied to sediment samples. Proxy types are ordered as they are discussed in the paper, with section numbers associated with each. Proxy types shown in grey can be used to reconstruct oxygen from benthic settings, and those in blue can be used for pelagic settings. Variations in thickness and opacity denote uncertainty and differing usage of terminology.</figcaption></figure>

be applicable to specific rock types or biogenic material or
may be interpreted in different ways depending on miner-
alogy or lithology. Programmes like the International Ocean
Discovery Program (IODP) have prioritized presenting litho-
logic metadata alongside formalized and accessible sample
archives. However, samples collected by individual laborato-
ries may not be associated with these data and/or archived in
an accessible way. Recently developed databases, such as the
Sedimentary Geochemistry and Paleoenvironments Project
(SGP), are working to circumvent some of these issues by
requiring lithologic context and detailed sediment descrip-
tions to accompany geochemical data submissions. Impor-
tantly, sedimentary features are crucial for guiding sample
selection for quantitative analyses, especially intervals that

are of interest because of specific redox characteristics. For
example, descriptors such as changes in organic carbon con-
tent, laminae, and the deposition of pyrites can be useful first
indicators of sedimentary redox/oxygen changes.

## 4.1 Historically based sedimentary redox/bottom water oxygen reconstructions

The presence/absence of laminae (example in Fig. 2) has
historically played an important role in reconstructing low-
oxygen systems, and it remains a popular tool today. The
presence of laminae is a key indicator of conditions that are
inconsistent with the survival of benthic fauna beyond sea-
sonal timescales, although microbioturbation of laminated
sediments, not visible to the naked eye, has been described
(Pike et al., 2001). Importantly, laminae can result from factors unrelated to redox changes and thus need to be interpreted with caution. For instance, laminated sediments are commonly found to be associated with diatom mats or giant diatoms, where diatom mats (e.g. *Thalassiothrix* spp.) suppress benthic activity (e.g. King et al., 1995; Kemp, 1996; Kemp et al., 2000, 2006; Grigorov et al., 2002). Laminae can also form due to grain size changes and particle sorting in sediment gravity flows, sediment-bed interaction, and seasonal to interannual changes in the grain size of settling particles (Kemp, 1996; O'Brien, 1996).

In addition to laminations, biofacies oxygen indices considering bioturbation, fauna, diversity, body size, and trophic levels have been used to characterize paleo-redox conditions, including specific oxygen levels (Rhoads and Morse, 1971; Behl and Kennett, 1996; Sperling et al., 2022). For instance, ichnological analysis has been widely applied to investigate ocean oxygenation (e.g. OAEs, glacial cycles, and hyperthermals) because different biofacies correspond to specific ranges of oxygen, such as anoxic, suboxic (low oxygen), and dysoxic (e.g. Casanova-Arenillas et al., 2022; Nicolo et al., 2010; Rodríguez-Tovar and Uchman, 2011; Rodríguez-Tovar, 2021). However, trace fossil occurrences could be impacted by both bottom water and pore water conditions. In some cases, trace fossils were produced during later favourable conditions (i.e. during diagenesis), and such traces are independent of the anoxic pore water conditions but are attributed to connections with favourable, more oxygenated bottom waters (e.g. Rodríguez-Tovar, 2021). Other environmental conditions, including food availability (e.g. organic carbon supply) and sedimentation rates, all need to be considered when interpreting ocean oxygenation (Rodríguez-Tovar, 2022). Recent work further demonstrates that carnivory and vision are linked to environmental oxygen levels (Sperling et al., 2013; McCormick et al., 2019). These indices have been used to reconstruct oxygen trends during the evolution of early animal life (Sperling et al., 2015; Canfield and Farquhar, 2009; Boyle et al., 2014; Tarhan et al., 2015; Van de Velde et al., 2018), oxygen impacts on mass extinctions (Reddin et al., 2020; Sampaio et al., 2021), and local oxygen levels independent of a broader evolutionary context.

The presence and relative abundance of pyrite and observations of its crystal structure can be indicators of water column euxinia (Fig. 2). Specifically, the size distribution of framboidal pyrite may reflect formation in euxinic versus more oxidizing water columns. Smaller framboids found in the Black Sea, for example, are interpreted to reflect a fast growth rate within the euxinic water column, as opposed to formation on longer timescales within sulfidic sediments (Wilkin et al., 1996, 1997; Wilkin and Barnes, 1997; Wignall and Newton, 1998). The size fraction of pyrite framboids has been applied within OAEs but also other intervals, with widespread geochemical evidence of marine anoxia (Wilkin et al., 1996; Wignall et al., 2005; Kuroda et

al., 2005; Jenkyns, 2010). Because syngenetic pyrite can incorporate or absorb trace elements (Huerta-Diaz and Morse, 1992), the trace element content of pyrite is also an important paleo-redox archive (Large et al., 2014).

## 4.2 Non-destructive methods for sediment analyses

Observations of sedimentary facies are used as a first-order evaluation of the depositional environment. Traditional methods (e.g. non-destructive core description and physical property measurements) are a fast, low-cost qualitative way to interpret redox/oxygen conditions. Over the past few decades, there have been important technological advances in describing sedimentary features in quantitative and non-destructive ways. Analytical instruments and imaging technology (e.g. microtomography, X-ray fluorescence (XRF) scanners, multi-sensor core loggers, scanning electron microscopy) can further improve the spatial and temporal resolution of the sedimentological and physical property measurements (see below), allowing for non-destructive and 3D detection of various sedimentary features on the sub-millimetre to micrometre scale.

X-ray computed tomography (CT) is a high-resolution (~0.1–1 mm) imaging technique that allows for visualization of the 3D structure of objects, determined by X-ray attenuation associated with variations in density and element compositions in sedimentary records (Fig. 2; see reviews by Mees et al., 2003; St-Onge et al., 2007). Standard-resolution CT imaging can be used on both whole round and section halves of sediment cores with minimal pretreatment. As a non-destructive method, it has been used to determine physical properties of sediments (e.g. density, porosity, and grain size; Fortin et al., 2013; Tanaka et al., 2011; Orsi et al., 1994; Amos et al., 1996; Mena et al., 2015) and to identify benthic communities (e.g. bioturbation analysis and trace fossil detection/ichnological analysis in the sediments, Dorador et al., 2020; Rodriguez-Tovar, 2022). Microstructure information obtained using standard-resolution CT has greatly improved the accuracy of sedimentological descriptions, whereas physical property data are critical for understanding oxygen penetration in the sediment profile and subsequent diagenetic processes controlled by pore water redox concentrations.

Multi-sensor core loggers (MSCLs) or multi-sensor tracks (MSTs) are widely used for continuous measurements of physical properties on centimetre scales in either whole round or section halves of sediment cores. These core loggers are usually equipped with detectors for measuring magnetic susceptibility, gamma ray density, natural gamma radiation, p-wave velocity, and resistivity, which provide density, porosity, and Fe-bearing mineral information for first-order evaluation of the ambient redox state in pore waters.

XRF scanners measure the relative abundance of elements (from Al to U, following the periodic table) on section half sediment cores at sub-millimetre to centimetre (i.e. high-resolution) scales in a non-destructive manner (Coudace et
al., 2006, 2019; Croudace and Rothwell, 2015). XRF data are considered to be semi-quantitative as elemental variability in the sediment cores is measured as counts and not concentrations. XRF data quality is affected by X-ray tube aging, water content, smoothness of the sample surface, and grain size (Böning et al., 2007; Tjallingii et al., 2007; Weltje and Tjallingii, 2008). Thus, appropriate sample preparation (e.g. core scraping and use of a thin polyester XRF film to smooth the surface) is required for high-quality data acquisition (Löwemark et al., 2019). Additionally, sediment composition (e.g. organic carbon and calcium carbonate content) may affect XRF counts because lighter elements (e.g. C, N, O) are outside of the XRF detection range. For instance, higher sedimentary organic carbon can dilute the number of counts for all elements (Löwemark et al., 2010). Normalization of the absolute counts with respect to an element that is less affected by biological and diagenetic processes (e.g. normalizing to Al or Ti) is used to assess the relative variability of elemental compositions (Löwemark et al., 2010) (see, for example, the Bering Strait in Fig. 2). Despite the limitations, scanning XRF is able to provide high-resolution data with fast and non-destructive measurements, allowing for a first-order assessment of redox-sensitive element abundance (e.g. Mo, U, Mn) prior to more labour-intensive analyses (e.g. solution-based bulk elemental concentration analyses, as discussed in Sect. 5).

# 5 Sedimentary redox trace elements and isotopes

## 5.1 Introduction

The potential for the concentrations of trace metals in sediments to act as proxies for past Earth surface conditions has been recognized since early observation of metal enrichments in organic-rich sediments (Goldschmidt, 1954). Trace metals provide some of the most commonly used proxies for the reconstruction of paleo-redox conditions in sediments (Algeo, 2004; Algeo and Maynard, 2004; Algeo and Rowe, 2012; Bennett and Canfield, 2020; Algeo and Li, 2020; Brumsack, 2006; Calvert and Pedersen, 1996; Little et al., 2015; Morford and Emerson, 1999; Nameroff et al., 2004; Scott and Lyons, 2012; Sweere et al., 2016; Tribovillard et al., 2006; Calvert and Pedersen, 2007; Zhou and McManus, 2023).

Sedimentary trace metal enrichments are associated with precipitation and/or adsorption of metals from ambient bottom and/or pore waters along a redox gradient (redox potential, Eh) primarily controlled by the decomposition of organic carbon using various oxidants (Calvert and Pedersen, 2007; Froelich et al., 1979). These redox reactions proceed in a well-defined sequence (Fig. 3), during which trace metals may be scavenged from ambient waters and subsequently enriched in sediments (i.e. authigenic enrichment as distinct from detrital input) as a result of changes in valence state

(e.g. Mn, U, Re, and Mo) and/or speciation (e.g. Cd, Ni, and Zn with solid-phase precipitation but no valence change) (Algeo and Li, 2020). Because sedimentary Eh varies in response to both bottom water oxygen availability and the rain rate of organic carbon, reconstructions using redox-sensitive elements to reconstruct bottom water oxygen must explicitly account for changes in the rain rate of organic carbon (see Sect. 5.4.1).

Trace element analysis has the advantage of facilitating “multi-proxy” data acquisition. Sedimentary trace metal concentration measurements are free from vital effects compared to trace metal incorporation into biogenic carbonate (e.g. foraminifera shells; see Sect. 8) and are particularly valuable when carbonate preservation is poor and sediments have remained relatively undisturbed post-deposition. Recently, a better understanding of redox-sensitive metal preservation in surface sediments and applications of statistical techniques have made it possible to quantify dissolved oxygen concentrations in coastal systems of the eastern Pacific (Valdés et al., 2021; Costa et al., 2023), paving the way for additional regional redox-sensitive trace metal calibrations and creating new possibilities for quantitative oxygen reconstructions.

In addition to redox-sensitive metal concentrations, isotope fractionation of these metals (e.g. Mo, U, Cr, and Fe) may occur during the exchange between seawater and other ocean sinks/sources (e.g. scavenging from reducing water columns) in various redox environments, making those isotope systems potential redox proxies. Technical advances in mass spectrometry have allowed for measurements of “non-traditional” stable metal isotope systems and have enabled their use in reconstructions of past ocean oxygenation changes (e.g. Andersen et al., 2017; Kendall et al., 2017; Severmann et al., 2008; Frei et al., 2011). Compared to authigenic enrichments, redox-sensitive metal isotope proxies may allow for more (semi-)quantitative redox reconstructions via isotope mass balance and potentially provide a more globally integrated perspective on ocean oxygen variability.

Trace element enrichments and their isotopes have provided key insights into ocean processes on various timescales (from the Precambrian to the present; Table 1), and research is ongoing to refine the interpretations of these proxies to shed new light on our understanding of global ocean oxygen responses to variations in Earth’s climate and other environmental variables.

## 5.2 Materials needed and analytical methods

Quantitative elemental concentrations are measured on dried sediments, which are either fully or partially dissolved (i.e. leached) to target authigenic phases. Samples are generally treated using bulk digestion methods such as acid digestion and alkaline fusion, depending on the sediment composition and elements of interest. To avoid contamination, all sample preparation should be performed in trace-metal-clean laboratories with acid-cleaned vessels and trace-metal-grade chem-
<figure><img src="image_4.png" /><figcaption>**Figure 2.** Examples of sedimentary features discussed in the main text. Left to right: clay to laminated gyttja sediments at the transition from the Baltic Ice Lake to the Littorina Sea during the Early Holocene of the Baltic Sea; laminae from the Bering Sea ODZ during the last deglacial; 3D CT scan from sediments deformed during the 2011 Tohoku Earthquake from offshore Japan; scanning XRF results (Si/Ti) from the Bering Sea; and a scanning electron microscope (SEM) image of framboidal pyrite from the Miocene Monterey Formation of California, USA (adapted after Berndmeyer et al., 2012; Kühn et al., 2014; Nakashima and Komatsubara, 2016).</figcaption></figure>

<table><caption>**Table 1.** Summary of redox trace elements and isotopes that can provide insights into ocean oxygenation from the Precambrian to the present.</caption><thead><tr><th>Proxy</th><th>Typical marine concentration of aqueous species</th><th>Marine isotope composition (± SD)</th><th>Ocean residence time (years)</th><th>Example reference</th></tr></thead><tbody><tr><td>Mn</td><td>~ 1.8 nmol kg<sup>-1</sup></td><td>NA</td><td>10–40</td><td>Bender et al. (1977)</td></tr><tr><td>Ce / Ce*</td><td>5 pmol kg<sup>-1</sup></td><td>N/A</td><td>50–130</td><td>Alibo and Nozaki (1999)</td></tr><tr><td>δ<sup>53</sup>Cr</td><td>4 nmol kg<sup>-1</sup></td><td>0.44–1.53‰ (?)</td><td>~ 3000</td><td>Qin and Wang (2017)</td></tr><tr><td>U and δ<sup>238</sup>U</td><td>13.4 nmol kg<sup>-1</sup></td><td>-0.39‰</td><td>400 000</td><td>Lau et al. (2019)</td></tr><tr><td>Re and δ<sup>187</sup>Re</td><td>~ 40 pmol kg<sup>-1</sup></td><td>-0.17 ± 0.12‰</td><td>130 000</td><td>Dickson et al. (2020)</td></tr><tr><td>V and δ<sup>51</sup>V</td><td>~ 35 nmol kg<sup>-1</sup></td><td>0.2 ± 0.07‰</td><td>~ 91 000</td><td>Nielsen (2020)</td></tr><tr><td>Mo δ<sup>98</sup>Mo</td><td>104 nmol kg<sup>-1</sup></td><td>2.34 ± 0.1‰</td><td>440 000</td><td>Kendall et al. (2017)</td></tr><tr><td>ε<sup>205</sup>Tl</td><td>65 nmol kg<sup>-1</sup></td><td>-6 ± 0.3 ε unit</td><td>~ 20 000</td><td>Owens (2019)</td></tr></tbody></table>

NA: not available.

Acid digestion and alkaline fusion are commonly used to dissolve sediment to analyse major, minor, and trace elements.

Major and minor element compositions can be measured using quantitative X-ray fluorescence (XRF), inductively coupled plasma optical emission/atomic emission spectroscopy (ICP-OES or ICP-AES), atomic absorption spectrophotometry (AAS), and microwave plasma atomic emission spectroscopy (MP-AES); minor and trace elements may be measured using inductively coupled plasma mass spectrometry (ICP-MS) that has a lower detection limit.

Quantification of redox-sensitive metal enrichment (Metal<sub>EF</sub>) may be determined following Tribovillard et al. (2006) (Eq. 1) and Böning et al. (2009) (Eq. 2):

$$ Metal_{EF} = (Metal/NE)_{sample} / (Metal/NE)_{background}, \quad (1) $$

where NE corresponds to the element for normalization.

$$ Metal \text{ excess (normalized by Al)} = Metal_{sample} - (Metal/Al)_{background} \times Al_{sample} \quad (2) $$

The upper continental crust has been widely used as a lithogenic background reference (Rudnick and Gao, 2003). Because lithogenic background ratios may vary by region, source materials (e.g. aeolian, river sediment input, and coastline or glacial erosion), and timescales, care should be taken to determine and cite an appropriate value (see Sect. 5.4.2).

Metal isotope measurements often target authigenic phases to avoid contamination from detrital components. As such, diluted acids (e.g. diluted HCl and HNO<sub>3</sub>) or weaker acids (e.g. acetic acid) are used in the partial digestion or leaching process. Leaching methods vary between and within
<figure><img src="image_5.png" /><figcaption>**Figure 3.** Redox ladder (modified from Fig. 1 in the Introduction) and redox-sensitive trace metals and metal isotopes discussed in this section. Dissolved oxygen ranges for OMZs, ODZs, anoxia, and dysoxia/hypoxia (low oxygen) are labelled in the figure. The redox ladder is modified from Canfield and Thamdrup (2009). The oxygen/redox potential range for the use of each redox-sensitive metal and metal isotope redox proxy is shown with the bars on the right.</figcaption></figure>

labs, even for the same isotope measurements (e.g. U isotopes, Tissot et al., 2018). Initial sample reconnaissance experiments should be used to determine the optimal leaching procedure. For high-precision stable metal isotope analysis, it is generally necessary to purify the element of interest from sample matrices to avoid possible spectral or non-spectral interferences on the instrument (e.g. through column chemistry).

Metal stable isotopes are now analysed routinely using thermal ionization mass spectrometry (TIMS) or multi-collector ICP-MS (MC-ICP-MS) instruments (Table 2). Precise and accurate stable isotope ratio measurements on either type of instrument depend on robust correction of instrumental mass bias produced during analysis (e.g. double-spike and sample-standard bracketing method) (e.g. Siebert et al., 2001; Ripperger et al., 2007; Tian and Wang, 2019; Nielsen et al., 2004, 2016; Wu et al., 2016). Additionally, metal stable isotopes can be measured by in situ techniques, including secondary ion MS (SIMS) and laser ablation MC-ICP-MS, which have shown unique potential in unravelling micrometre-scale information from samples with complex textures or zonation that are otherwise inaccessible by bulk

analysis. Currently, in situ stable isotope analysis is more frequently used in studies of high-temperature and cosmo-genic processes, as well as environmental conditions of early Earth. This leaves ample opportunities to adapt existing in situ methodologies and develop new ones for more recent paleoceanographic research.

## 5.3 Redox-sensitive metal and metal isotope proxies

### 5.3.1 Redox-sensitive metal proxies with valence state changes by redox potential

#### Manganese

Manganese (Mn) has three oxidation states (II, III, and IV). The reduced forms of Mn are soluble in low-oxygen waters ($<10 \mu mol \text{ kg}^{-1} \text{O}_2$) (Madison et al., 2013; Oldham et al., 2017), which include Mn(II) and soluble Mn(III) complexed by inorganic or organic ligands (Mn(III)-L) (Oldham et al., 2015). The oxidized form of Mn(IV) forms solid Mn(IV) oxides. Consequently, the residence time of dissolved Mn in the oxygenated deep ocean is of the order of 10–40 years (Bender et al., 1977; Klinkhammer and Bender, 1980; Hayes et al., 2018). As reduced Mn(II) can be oxidized to Mn(III) / Mn(IV) oxyhydroxides with even micromolar levels of oxygen (Tebo et al., 2004; Morgan, 2005; Clement et al., 2009), sedimentary Mn enrichment can be used as an oxic indicator in pore waters (Burdige and Gieskes, 1983; Froelich et al., 1979; Calvert and Pedersen, 1996). However, free Mn(II) can also precipitate as $MnCO_3$ and/or co-precipitate with authigenic calcite in reducing pore waters with high alkalinity, for example, when methanogenesis occurs (Calvert and Pedersen, 1996; Mucci, 2004), which may lead to a false positive for oxic conditions. Thus, Mn should be evaluated simultaneously with other redox-sensitive metals.

#### Iron

The iron paleo-redox proxy can be used to distinguish oxic, ferruginous, and euxinic water column settings (reviewed in Raiswell et al., 2018). In oxic environments, Fe exists as Fe oxyhydr(oxides), including ferrihydrite, lepidocrocite ($\gamma$-FeOOH), goethite ($\alpha$-FeOOH), hematite ($\alpha$-Fe$_2$O$_3$), maghemite ($\gamma$-Fe$_2$O$_3$), and magnetite (Fe$_3$O$_4$). As the ambient seawater becomes depleted in oxygen, Fe oxyhydr(oxides) can be reduced to Fe(II). With sulfide production during sulfate reduction, reduced Fe(II) can be converted to Fe sulfides that include mackinawite (FeS), greigite (Fe$_3$S$_4$), and pyrite (FeS$_2$). In strongly reducing waters (e.g. methanic conditions), siderite ($FeCO_3$) can also form. Combined with Fe sulfides, carbonate-bearing Fe and Fe (oxyhydr)oxides make up the highly reactive Fe pool because these forms of Fe readily react with free sulfide (e.g. $HS^-$) in early diagenetic stages. By leaching out different Fe phases, Fe speciation uses highly reactive
<table><caption>**Table 2.** Quantitative analytical methods for trace metals and metal isotopes.</caption><thead><tr><th>Analysis</th><th>Digestion method</th><th>Instrument</th><th>Quality control</th></tr></thead><tbody><tr><td>Bulk major elements</td><td>Full digestion (alkaline fusion or acid digestion)</td><td>ICP-OES (rapid and cost-efficient), solution-based or laser ablation (for LiBO<sub>2</sub>-fused beads) ICP-MS (low detection limit), XRF, AAS</td><td>Instrumental drift correction (e.g. internal standards) and standard reference materials</td></tr><tr><td>Bulk minor trace elements</td><td>Full digestion (alkaline fusion or acid digestion)</td><td>Usually ICP-MS, solution based or laser ablation of LiBO<sub>2</sub>-fused beads</td><td>Instrumental drift correction (e.g. internal standards) and standard reference materials</td></tr><tr><td>Stable metal isotopes</td><td>Leaching authigenic phases</td><td>TIMS, (laser ablation) MC-ICP-MS, and in situ SIMS</td><td>Double-spike and sample-standard bracketing</td></tr></tbody></table>

Fe (Fe<sub>HR</sub>) / total Fe (Fe<sub>T</sub>) and pyrite Fe (Fe<sub>py</sub>) / Fe<sub>HR</sub> to distinguish between oxic, ferruginous, and sulfidic conditions. Modern sediment calibrations indicate a threshold of Fe<sub>HR</sub> / Fe<sub>T</sub> > 0.38 for anoxic water columns. Under anoxic regimes (Fe<sub>HR</sub> / Fe<sub>T</sub> > 0.38), Fe<sub>py</sub> / Fe<sub>HR</sub> has been used to differentiate sulfidic (Fe<sub>py</sub> / Fe<sub>HR</sub> > 0.7 ~ 0.8) from ferruginous (Fe<sub>py</sub> / Fe<sub>HR</sub> < 0.7) waters. When Fe<sub>HR</sub> / Fe<sub>T</sub> < 0.38, high Fe<sub>py</sub> / Fe<sub>HR</sub> values (> 0.8) have also been used to indicate oxic water columns with pore water sulfide accumulation in organic-rich sediments.

## Uranium, rhenium, and vanadium

Uranium (U), rhenium (Re), and vanadium (V) behave conservatively in seawater – the residence time in the ocean is ~750 000 years for Re (Akintomide et al., 2021), ~300 000–600 000 years for U (Dunk et al., 2002; Ku et al., 1977; McManus et al., 2005; Morford and Emerson, 1999; Lau et al., 2019), and ~50 000–100 000 years for V (Shiller and Boyle, 1987; Tribovillard et al., 2006; Nielsen, 2020). As a result, sedimentary concentration changes in U, Re, and V on timescales shorter than tens of thousands of years are likely not a response to the changes in the dissolved concentration in the overlying water column. Instead, the downward flux of metal reduction, in accordance with the redox potential of the pore water, is likely the driver of the sedimentary variations (Böning et al., 2004; Colodner et al., 1995; Sundby et al., 2004), making these elements potentially useful oxygen indicators.

Rhenium exists as ReO<sub>4</sub><sup>-</sup> in oxic waters but can be reduced to Re(IV) oxides (e.g. ReO<sub>2</sub>) in reducing environments. The redox potential of the Re(VII)-Re(IV) couple is higher than that of U(VI)-U(IV), situated between MnO<sub>2</sub>-Mn(II) (manganous) and Fe<sup>3+</sup>-Fe<sup>2+</sup> (ferruginous), and is similar to the redox potential of NO<sub>3</sub><sup>-</sup>-NO<sub>2</sub><sup>-</sup> (Bratsch, 1989; Algeo and Li, 2020). Re preservation in sediment could also be associated with thiolation of ReO<sub>4</sub><sup>-</sup> to particle-reactive ReO<sub>n</sub>S<sub>4-n</sub>, which enhances its particle reactivity towards iron sulfides (Akintomide et al., 2021) and/or co-precipitation with the Fe-Mo-S phase in sulfidic waters (waters in which oxygen

is undetected and sulfide is present) (Helz and Dolor, 2012; Helz, 2022). Free sulfide levels in the most sulfidic water columns are still insufficient to support thiolated ReO<sub>4</sub><sup>-</sup> as a major species due to higher required sulfide levels relative to molybdate thiolation (Helz and Dolor, 2012; Vorlicek et al., 2015). However, this potential exists in euoxic pore waters (Akintomide et al., 2021). Less is known about Re isotopes and their usefulness for constraining past changes in ocean oxygenation, although early studies have worked on constraining the Re isotope mass balance (Dellinger et al., 2021; Dickson et al., 2020).

Vanadium mainly occurs as V(V) in oxic waters in the form of vanadate (e.g. HVO<sub>4</sub><sup>2-</sup> and H<sub>2</sub>VO<sub>4</sub><sup>-</sup>). However, unlike U and Re, vanadate can be scavenged by adsorption onto Fe-Mn (oxyhydr)oxides and clay minerals (e.g. Wehrli and Stumm, 1989; Morford and Emerson, 1999). The redox potential of the V(V)-V(IV) couple is similar to that of the Re(VII)-Re(IV) couple (Algeo and Li, 2020). Thus, as oxygen draws down, vanadate can be reduced to the V(IV) species (vanadyl, VO<sup>2+</sup>, and VO(OH)<sup>3+</sup>) by organic compounds, which can co-precipitate/complex with mineral particles and organic matter (Emerson and Huested, 1991; Algeo and Maynard, 2004). Under more reducing conditions (e.g. sulfidic), vanadyl might be further reduced to the V(III) species by free sulfide in the ambient waters, which precipitate as solid oxides (V<sub>2</sub>O<sub>3</sub>) or hydroxides (VOOH) (Wanty and Goldhaber, 1992). Despite the different authigenic enrichment mechanisms, V reduction and sequestration into sediments still begin under low-oxygen conditions, making it a tracer of such conditions.

In oxic water columns, U exists as soluble U(VI) and binds to carbonate ions, forming Ca<sub>2</sub>UO<sub>2</sub>(CO<sub>3</sub>)<sub>3</sub> (Endrizzi and Rao, 2014, Langmuir, 1978). The redox potential of the U(VI)-U(IV) couple is below that of the Fe<sup>3+</sup>-Fe<sup>2+</sup> couple but above SO<sub>4</sub><sup>2-</sup>-H<sub>2</sub>S (Fig. 1, Morford and Emerson, 1999; Zheng et al., 2002a, b). In reducing environments, U(VI) turns into U(IV) in the form of solid uraninite (UO<sub>2</sub>) or adsorbs onto sediment solids, which may involve biologically mediated processes (Crusius et al., 1996; Klinkhammer and Palmer, 1991; McManus et al., 2005; Zheng et al., 2002a,
b; Lovley et al., 1991; McManus et al., 2006; Stirling et al., 2015; Rolison et al., 2017).

As discussed above, soluble U(VI) can be reduced to insoluble U(IV). The sedimentary U sequestration process also introduces significant isotopic fractionation (e.g. Zhang et al., 2020), as the nuclear volume effect causes a preferential removal of the heavy $^{238}$U relative to the lighter $^{235}$U isotope. Because of the long residence times of U (~300 000–600 000 years; Dunk et al., 2002), the isotopic composition of U in seawater is globally homogenous (–0.39‰ in the modern ocean; Andersen et al., 2017). Uranium uptake in reducing sediments is the primary U sink in the global ocean, and, hence, seawater $\delta^{238}$U changes can be associated with the extent of seafloor anoxia (e.g. Lau et al., 2019, 2020).

Oxic sediment deposits that record the seawater U isotope value (e.g. shallow marine carbonates and Mn oxide crusts) have been used to infer the areal extent of anoxic sinks in the global ocean using isotope mass balance models (e.g. Zhang et al., 2018, 2020). However, post-depositional diagenesis of carbonate could result in much larger offsets from the seawater U isotope value (e.g. Chen et al., 2022). In contrast, sediments deposited within anoxic conditions, such as organic-rich black shales, will typically record enriched $\delta^{238}$U values during more intense anoxia, although the expression of isotope enrichment is complicated by processes that vary across depositional environments (e.g. diffusion of U between the sediment-water interface and the zone of U reduction within the sediment; see Andersen et al., 2014). Long-lasting anoxia/euxinia within restricted basins (e.g. limited water renewal) is shown to have a larger fractionation factor from seawater (~+0.5‰ to +0.7‰ based on modern observation) due to diffusion limitations of U (VI) in shallow sediments, where most U reduction occurs (e.g. Lau et al., 2020). Within anoxic facies, carbonate-associated uranium isotopes have also been used to infer local deoxygenation in sediments, with the large advantage that this proxy is not significantly impacted by post-depositional oxidation (Clarkson et al., 2021b).

## Molybdenum

In oxic waters, Mo primarily exists as soluble molybdate ($Mo(VI)O_4^{2-}$) and behaves conservatively (~440 000-year residence time) (Miller et al., 2011). However, ~30%–50% of molybdate may be sequestered through adsorption onto Mn and Fe (oxyhydr)oxides in oxic waters (Kendall et al., 2017). Unlike the low-oxygen indicators (e.g. U, Re, and V), reductive Mo removal requires sulfidic conditions that lead to progressive thiolation of molybdate (thiomolybdate series $Mo(VI)O_xS_{4-x}^{2-}$, x = 0~4) (Hlohowskyj et al., 2021). Thiomolybdates are particle reactive and readily scavenged into sediments and onto iron sulfides (Freund et al., 2016). Mo removal from sulfidic pore waters has been associated with a hydrogen sulfide ($H_2S$) threshold of > 11 µmol (when $MoS_4^{2-}$ starts to dominate in the waters, Helz et al., 1996).

Yet, recent studies have suggested that under weakly sulfidic conditions ($[H_2S] < 11 µM$), intermediate thiomolybdate species could be the dominant Mo species in the water column that contribute to Mo sequestration (e.g. Tessin et al., 2018). Multiple pathways have been proposed for Mo removal from sulfidic waters, including (1) the organic matter (OM) pathway that leads to Mo (IV or VI)-OM complexes (Dahl et al., 2017), (2) the Fe–sulfide pathway that has thiolated Mo adsorption to iron sulfide phases with subsequent Mo(VI) reduction to Mo(IV) (Miller et al., 2020) and/or that incorporates Mo(IV) into Mo–Fe–S structures such as $FeMoS_2(S_2)$ (Helz and Vorlicek, 2019), and (3) the biological pathway that implies biological uptake (e.g. by sulfate-reducing bacteria) and Mo reduction by enzymes (e.g. Dahl et al., 2017). Authigenic Mo enrichment has thus been interpreted as an indicator of sulfidic environments provided that coeval enrichments of other redox-sensitive trace metals (e.g. U or Re that is not scavenged by Mn oxides in oxic environments) are observed.

As discussed above, in sulfidic environments Mo from oxic waters (mostly as $MoO_4^{2-}$) is converted into particle-reactive thiomolybdate species in the presence of free sulfide. During this transformation the Mo isotopes are fractionated, with the more sulfidized thiomolybdate species becoming isotopically lighter relative to seawater (He et al., 2022; Kerl et al., 2017; Tossell, 2005). The Mo isotopic composition of mildly euxinic sediments is thus expected to be lighter than the seawater value (~2.34‰ in the modern ocean; Nakagawa et al., 2012; Nägler et al., 2013). Low $\delta^{98}$Mo values of sediments deposited in such environments can be modelled assuming higher scavenging rates for the more sulfidized Mo species, allowing for semi-quantitative reconstructions of $H_2S$ concentrations (Dahl et al., 2010a; Matthews et al., 2017; Sweere et al., 2021). However, in strongly sulfidic conditions (e.g. the Black Sea), the conversion of $MoO_4^{2-}$ to $MoS_4^{2-}$ (the most sulfidized species) is nearly complete (Erickson and Helz, 2000) such that little to no fractionation is expressed (Tossell, 2005). The sedimentary $\delta^{98}$Mo values therefore approach seawater compositions (Neubert et al., 2008; Brüske et al., 2020).

In conclusion, sedimentary Mo isotopes can be used to trace the local euxinic water column conditions when the global seawater Mo isotopic composition is known or close to modern values (e.g. Holocene/Pleistocene to Paleocene sediments; Andersen et al., 2018; Azrieli-Tal et al., 2014; Hardisty et al., 2021b; Matthews et al., 2017; Sweere et al., 2021; Riedinger et al., 2021). Environments with a quantitative drawdown of dissolved Mo may also be studied to infer global seawater $\delta^{98}$Mo values to estimate the global-scale extent of oxic–anoxic–euxinic Mo sinks on geological timescales (e.g. Dahl et al., 2021; Dickson and Cohen, 2012; Dickson et al., 2016). Sedimentary (co)variations in $\delta^{98}$Mo and $\delta^{238}$U have also been applied to trace past changes in ocean oxygenation, especially on orbital to million-year
timescales (e.g. Andersen et al., 2020; Chen et al., 2015; Chiu et al., 2022; Clarkson et al., 2021a; Dahl et al., 2010a; Dickson, 2017; Gordon et al., 2009; Hardisty et al., 2021b; Kendall et al., 2015; Sweere et al., 2021; Wang et al., 2016a; Zhang et al., 2018).

### 5.3.2 Trace metal proxies with speciation changes in sulfidic waters (cadmium, zinc, and nickel)

Unlike the previously discussed metals, Cd, Zn, and Ni behave like micronutrients, with a depletion in the surface ocean due to biological uptake and increasing concentrations at depth due to decomposition of sinking organic material (Bruland, 1983; Flegal et al., 1995; Nozaki, 1997; Bruland and Lohan, 2003). Authigenic metal enrichments in the sediments primarily occur in sulfidic waters because they can form insoluble sulfides (e.g. Cd and Zn; Rosenthal et al., 1995; Tribovillard et al., 2006; Little et al., 2015) and/or be incorporated into the pyrite structure (e.g. Huerta-Diaz and Morse, 1992; Large et al., 2014), making them a possible tracer of sulfidic conditions. However, we note that Ni has also been linked to sinking fluxes of organic material in upwelling regions because of its close link with productivity and less diagenetic alteration associated with sedimentary sulfur and manganese cycling (Böning et al., 2015).

## 5.4 Factors controlling trace metal preservation/metal isotope fractionation

In addition to the redox potential and organic carbon rain rate, sedimentary enrichments of trace metals can be altered by several processes, which may affect delivery to the sediments and result in remobilization/recycling in the sediment pore waters. These caveats may lead to decoupled sedimentary responses from the water column oxygen variability.

Due to anthropogenic activities, coastal marine ecosystems are susceptible to pollution by potential contaminant metals and metalloids from industrial and domestic waste. Some pollutant metals that have adverse impacts on aquatic life and human health are also redox sensitive, such as Mo, Zn, B, Mn, Ba, Co, Ni, Sr, Cr, Cd, Zr, V, Cu, and Ce. In some regions, the concentrations of these elements have recently increased, and the values do not align with increases in lithogenic background inputs (Muñoz et al., 2022; Valdés et al., 2023). As river sediments may concentrate and deliver anthropogenically sourced metals during transport to the ocean (Pizarro et al., 2010; Yevenes et al., 2018), using standard lithogenic background corrections may overestimate enrichment factors as the authigenic fraction will be augmented with the anthropogenic input. In regions with high anthropogenic input, it is thus recommended to use the local lithogenic background in the region of study (e.g. crust, river, dust, wetland sediment) (Muñoz et al., 2022, 2023; Valdés and Tapia, 2019; Valdés et al., 2023).

### 5.4.1 Organic carbon and trace element burial

Because redox potentials vary as a function of both reducible and oxidant availability, enrichments of trace elements in sediments could result from bottom water oxygen availability and/or the rain rate of organic material. For example, glacial authigenic U enrichments in the Southern Ocean have been found to occur primarily as a function of changes in export production (e.g. Chase et al., 2001; Kumar et al., 1995). As some redox-sensitive metals (e.g. Mo, V, Ni, and Cd) can be concentrated in plankton due to biological uptake, they are also efficiently transported to the sediments via biodtriatus and particulate organic matter (e.g. in upwelling areas) (Böning et al., 2004; Muñoz et al., 2012; Valdés et al., 2014; Castillo et al., 2019; Nameroff et al., 2004; Muñoz et al., 2023). Thus, to isolate bottom water oxygen concentrations, reconstructions using redox-sensitive trace elements must be accompanied by independent constraints on the supply of particulate organic carbon (e.g. Anderson et al., 2019; Bradtmiller et al., 2010; Jaccard et al., 2009; Jacobel et al., 2020; Pavia et al., 2021). It is particularly important that proxies for organic carbon flux are independent of oxygen, rendering classic proxies like total organic carbon (TOC) ineffective since its sedimentary abundance is itself a function of oxic respiration (e.g. Burdige, 2007; Tyson, 2001). For a description of marine organic compounds that are susceptible to oxygen, see Sect. 6 on organic biomarkers as proxies for seawater oxygen.

Trace element enrichment proceeds according to sequential redox thresholds, and attempts have been made to define bottom water oxygen thresholds, below which redox-sensitive trace elements would be expected to become enriched in the sediments (e.g. Algeo and Li, 2020; Bennett and Canfield, 2020). Unfortunately, this approach is inappropriate for reconstructing bottom water oxygen concentrations because variations in organic carbon can be the primary determinant of trace element enrichment.

### 5.4.2 Detrital influences on authigenic enrichments of trace metals

Shifts in sedimentary elemental compositions may be associated with changing proportions of sediment sources (e.g. lithogenous, biogenous, and hydrogenous) with inherently different elemental matrices. For example, nearshore sediments are highly influenced by terrestrial inputs (e.g. fluvial and aeolian sediments) and organic fluxes from primary productivity, whereas deep-sea sediments generally receive only the finest fraction of lithogenic particles (e.g. clays from dust). Estimates of authigenic concentrations are based on an assessment of the lithogenic contribution (detrital) using Al as the most common approximation. This is based on the conservative behaviour of Al during weathering and soil formation and the assumption that Al concentrations are very similar in most common sedimentary rocks (Calvert and Ped
ersen, 2007), which may not hold true on a global scale. Ti-
tanium has also been used for normalization for lithogenic
contribution, but the concentration of this element is more
variable than Al in different rock types (Calvert and Ped-
ersen, 2007). Additionally, the estimate of detrital contribu-
tions assumes that the detrital element (e.g. Al) analysed is
only in the aluminosilicate fraction. Therefore, estimations
attributed to other phases (e.g. hydrogenous) could be under-
estimated (Van der Weijden, 2002). Caution should be em-
ployed when correlating normalized data because it could
modify the original correlations between elements (Cole et
al., 2017).

**5.4.3 Post-depositional diagenetic effects on sedimentary trace metals**

As with many other proxies, the utility of redox-sensitive
metals as a paleo-oxygen indicator relies on the post-
depositional persistence (or preservation) of the initial redox
signal. Modifications to trace element distributions during di-
agenesis include whole-scale overprinting of sedimentary re-
dox signals (e.g. Jacobel et al., 2020; Zheng et al., 2002a, b),
partial removal (e.g. Bonatti et al., 1971; Chase et al., 2001;
Jacobel et al., 2017; Morford et al., 2009), and oxidation and
downcore precipitation along concentration gradients (Col-
ley et al., 1989; Colley and Thomson, 1985; Jacobel et al.,
2017; McKay and Pedersen, 2014; Anschutz et al., 2002; De-
flandre et al., 2002).

One of the most significant pore water alterations that
can modify the original sedimentary metal enrichment signal
is post-depositional organic matter remineralization, which
progressively consumes pore water oxygen and changes lo-
cal redox potential (Morford and Emerson, 1999; Nameroff
et al., 2002) (Fig. 1). As oxygen is depleted, Fe / Mn (oxy-
hydr)oxides are reduced to release Mn(II) and Fe(II) that dif-
fuse into the pore waters. Aqueous Mn(II) may diffuse up-
wards until it reaches oxygenated pore waters and can re-
precipitate as Mn oxides (Lynn and Bonatti, 1965; Burdige
and Gieskes, 1983). As a result, a post-depositional Mn spike
may occur right above the depth where the pore water oxy-
gen concentration goes to zero (Burdige and Gieskes, 1983;
Froelich et al., 1979). Preservation of this peak is affected by
subsequent variability in the oxygen penetration depth (e.g.
Finney et al., 1988; Mangini et al., 2001; Anschutz et al.,
2002; Deflandre et al., 2002). Shoaling of the oxygen pen-
ation depth would push the Mn peak into reducing pore
waters, in which it will dissolve and diffuse upwards, leaving
no trace of the former peak (Froelich et al., 1979). In contrast,
if the oxygen penetration depth increases in the sediment, the
Mn peak will be preserved because it will remain within the
oxygen-rich zone (e.g. Froelich et al., 1979; Mangini et al.,
1990; Deflandre et al., 2002). Therefore, it has been proposed
that peaks in authigenic Mn concentrations in sediments are
best interpreted as pore water oxygen concentrations increas-

ing over time rather than indicators of (static) high oxygen
concentrations (e.g. Volz et al., 2020; Pavia et al., 2021).

Reductive dissolution of Fe / Mn (oxyhydr)oxides may lead to additional metal release as they are carrier phases for many trace metals (Algeo and Tribovillard, 2009; Scholz et al., 2011, 2017). For instance, remobilized V due to Mn oxide reductive dissolution (Seralathan and Hartmann, 1986; Legeleux et al., 1994; Hastings et al., 1996) may either diffuse upwards into bottom waters (Heggie et al., 1986; Shaw et al., 1990) or diffuse downwards and re-precipitate at a deeper sediment depth (Colley et al., 1984; Jarvis and Higgs, 1987). Post-depositional build-up of reducing conditions (e.g. sulfate reduction) would also facilitate additional trace metal sequestration (e.g. Mo) by sedimentary organic material and/or sulfides (e.g. pyrite or other metal sulfides) (Al-Farawati and van den Berg, 1999; Erickson and Helz, 2000; Helz et al., 1996; Helz and Vorlicek, 2019).

Remobilization of authigenic U has also been studied ex-
tensively in regions characterized by large oscillations in
pore water redox potential (e.g. Morford et al., 2009). When
pore water oxygen increases, U remobilization may occur
and allow U diffusion to the overlying bottom water and/or
re-precipitation at a deeper depth (Fig. 4).

These observations have several important implications. An absence of redox-sensitive metal enrichment cannot be taken as evidence that such conditions were absent, as short-lived events may not be recorded. This is especially true in environments with accumulation rates less than 2 (Jung et al., 1997; Mangini et al., 2001) or 3 cm kyr⁻¹ (Jacobel et al., 2020), where pore waters may retain active redox fronts long after the time of initial deposition, especially if sedimentary organic carbon is low. Caution is also needed in interpreting the shape of sedimentary enrichment features as primary signals, and both the sharpness of peaks and their temporal structure (Crusius and Thomson, 2000; Jacobel et al., 2020, 2017; Thomson et al., 2000) may be modified post-depositionally.

**5.4.4 Sedimentation rate changes**

The impact of sedimentation rate on authigenic enrichment
should also be considered when evaluating metal accumu-
lation. Changes in sedimentation rate would be expected to
impact metal accumulation since sedimentation rate directly
influences the rate of organic carbon respiration and depth of
bioturbation. Sedimentation rate has been used as a proxy for
the flux of organic carbon to the sediment-water interface,
with higher rates associated with more reducing conditions
and a shoaling of the oxygen penetration depth (Boudreau,
1994; Tromp et al., 1995). A shallower oxygen penetration
depth would reduce pore water exposure to oxygen and al-
low for a better preservation of trace metals (Fig. 5). A
special case is the occurrence of instantaneous depositional
events (e.g. turbidite layers), which could introduce pulses of
sediment delivery that significantly reduce oxygen exposure
<figure><img src="image_6.png" /></figure>

**Figure 4.** Schematic of authigenic U (aU) post-depositional diagenesis, after Jacobel et al. (2020). (a) In a baseline scenario, the relatively low-oxygen concentration at a certain depth below seafloor (blue line) leads to aU precipitation (green line). (b) As the bottom water oxygen concentration and the depth of oxygen penetration both increase, a portion of the previously precipitated aU becomes remobilized and diffuses upwards (green arrow). The rest of the remobilized U re-precipitates downcore. (c) Bioturbation (black arrows) may also mix aU-containing sediment upwards, exposing it to a better-oxygenated environment, where aU may be remobilized and released back into the bottom water (Morford et al., 2009).

of the underlying sediments. Rapid sediment accumulation would then facilitate build-up of reducing pore waters that lead to diagenesis (e.g. Fe / Mn reduction) (Anschutz et al., 2002; McKay and Pedersen, 2014; Wang et al., 2019a). However, an increase in the rate of non-reactive sediment accumulation can also dilute the relative concentrations of organic matter and trace metals while reducing the downward diffusion of dissolved gases (oxygen) or aqueous species (trace metals, sulfate). Modelled authigenic Mo and U as a function of sedimentation rate show dramatic decreases in their authigenic concentration with increasing sedimentation rate (Liu and Algeo, 2020; Hardisty et al., 2018; Morford et al., 2007).

### 5.4.5 Particulate shuttle and basin effect on trace metal delivery to the sediments

In many cases, trace metal enrichments may be controlled by a “particulate shuttle” (e.g. particulate Fe / Mn (oxyhydr)oxides and phytoplankton remains, Zheng et al., 2002a, b; Algeo and Tribovillard, 2009; Tribovillard et al., 2012; Sweere et al., 2016; Scholz et al., 2017; Ho et al., 2018; Severmann et al., 2008; Muñoz et al., 2023) and the re-supply of certain trace metals (“basin reservoir effect”; e.g. Algeo and Lyons et al., 2006). Due to affinity of Mo on Fe / Mn (oxyhydr)oxides, molybdate adsorbs onto particulate Fe / Mn oxyhydroxides in the oxic waters while being transferred through the water column. These particles are then reduced when oxygen is depleted in the ambient waters, releasing molybdate that either diffuses back into the water column or is scavenged in the sulfidic sediments (Morford and Emerson, 1999; Morford et al., 2005). The latter process accelerates the transfer of authigenic Mo to the sediment relative to other redox-sensitive trace metals (e.g. U) that are not affected by the particulate shuttle, leading to el-

evated authigenic Mo / U ratios in the sediments (e.g. Caraco Basin, Algeo and Tribovillard, 2009). Operation of particulate Fe–Mn (oxyhydr)oxide shuttles occurring close to a (variable) redoxcline in the water column could thus be interpreted from Mo–U covariation in the sediments.

On the other hand, dissolved trace element supply is limited in hydrographically restricted basins compared to the open ocean. Subsequent scavenging of certain trace metals may deplete the dissolved metal reservoir in the water column. Consequently, trace metal enrichment can vary considerably in restricted ocean areas compared to an open ocean setting (Algeo and Lyons, 2006; Algeo and Rowe, 2012; Sweere et al., 2016; Algeo and Li, 2020; Bennett and Canfield, 2020). For instance, dissolved Mo / U in the Black Sea is lower than that in the open ocean due to continuous scavenging of Mo in the water column, leading to a lower authigenic (Mo / U) ratio (Algeo and Tribovillard, 2009). Redox-sensitive enrichments may thus vary significantly in different depositional systems depending on metal delivery from the water column. The relationship between Mo and U enrichment factors (MoEF and UEF) to establish the shuttle effect and to infer the oxygenation conditions of the depositional environment can be found from the model proposed by Algeo and Tribovillard (2009) and Tribovillard et al. (2012).

It is necessary to take site-specific impacts and changes in environmental variables through time, such as organic carbon rain rate, into account when linking trace metal enrichments to redox conditions for each individual depositional system (Algeo and Li, 2020).

### 5.4.6 Interpretive approaches for reducing uncertainty

Several studies have applied the use of a suite of elements, including ratios and corrections for detrital phases, to compensate for some of the issues mentioned above, exploiting
<figure><img src="image_7.png" /><figcaption>**Figure 5.** Schematic of aU preservation in sediments with varying sedimentation rates, after Costa et al. (2018). (a) In a baseline scenario, the relatively low-oxygen concentration at a certain depth below seafloor (blue line) leads to aU precipitation (green line). (b) When sedimentation rates are low (the dashed blue line is the original seafloor position), as the bottom water oxygen concentration increases, aU remobilization takes place, and preservation is poor. (c) When sedimentation rates are high, aU is insulated from the oxygen penetration down the sediments, and the aU signal is better preserved.</figcaption></figure>

the variable response of different elements to the array of controlling parameters (Algeo and Lyons, 2006; Crusius et al., 1996; Jones and Manning, 1994). A recent review has found some redox dependency for all studied trace element ratios but also stressed several complications, including the need for these proxies to be carefully calibrated for each individual setting (Algeo and Li, 2020).

With an understanding of individual preservation mechanisms, the use of multiple redox proxies leads to a more nuanced interpretation of past conditions. Chromium, Re, and Mo have been used to discern the global extent of anoxic and sulfidic conditions (Reinhard et al., 2014). Molybdenum and U have similarly been used to distinguish between sulfidic or non-euxinic conditions and changes over time while also providing evidence of water mass restriction (Zhang et al., 2022; Algeo and Tribovillard, 2009). Although it is possible that Re and Mo co-precipitate, varying Re/Mo ratios might provide evidence of dissolution of other carriers, thereby increasing the delivery of Re or Mo to pore waters (Helz, 2022). For instance, in the Humboldt Upwelling Ecosystem (northern Chile and central Peru), Mo/U and Re/Mo ratios have been used to differentiate suboxic conditions (low oxygen and non-euxinic) from anoxic conditions in the depositional environment (Salvatteci et al., 2014, 2016; Valdés et al., 2014, 2021; Castillo et al., 2017). Higher Re/Mo and lower Mo/U than the seawater value (Crusius et al., 1996) would reflect suboxic (low-oxygen) conditions in the absence of H₂S (Tribovillard et al., 2006; Algeo and Tribovillard, 2009). On the contrary, in the presence of reducing and occasionally sulfidic bottom waters, Mo accumulation increases relative to U, implying that the Mo/U ratio in sediments could be equal to or higher than that of the water column

(molar ratio of 7.53 ± 0.25, Millero, 1996). However, the fidelity of these interpretations hinges on a clear understanding of differing authigenic preservation mechanisms, redox or non-redox related (e.g. organic carbon flux, particulate shuttle and basin reservoir effect), and the potential for diagenetic loss that would compromise the record.

Other empirical geochemistry proxies have also been proposed to evaluate depositional settings (Algeo and Li, 2020, and references therein) (e.g. restricted basins vs. continental margins with intensive upwelling). For instance, a decrease in sedimentary Mo/TOC ratios has been associated with water mass restriction in anoxic marine environments (e.g. in a silled basin) based on the observation that water column Mo can be depleted in stagnant basins (Algeo and Lyons, 2006). Based on the close association of Cd and productivity (Sect. 5.3.2, Horner et al., 2021), elevated Cd/Mo ratios may be used to indicate upwelling zones on the continental margin. Low sedimentary Cd/Mo (close to seawater value) caused by metal depletion in sulfidic water columns and high Co/Mn values attributed to a dominated river supply over a deep metal source (e.g. from upwelling) have been used to indicate restricted basin environments (Sweere et al., 2016). We caution that these empirical relationships require a better mechanistic understanding for trace metal cycling and can only be used in the marine environments where they have been calibrated.
## 5.5 Future perspectives

### 5.5.1 Towards quantitative oxygen proxies on a local scale

A few recent studies have investigated the potential for using redox-sensitive trace metals in quantitative oxygen reconstructions, especially on a local scale where other contributing factors to metal enrichments are less variable (e.g. sedimentation rates and lithogenic background). For instance, Costa et al. (2023) develop U/Ba as a local- to regional-specific bottom water oxygen proxy, which explicitly takes organic carbon rain rate into account via normalization with respect to Ba. Local U/Ba calibrations for the Arabian Sea, eastern equatorial Pacific (EEP), and western equatorial Pacific suggest that U/Ba may be used to capture bottom water oxygen concentrations in regions with > 50 µmol kg⁻¹ oxygen and high oxygen variability (several tens of µmol kg⁻¹). This shows the potential for redox-sensitive trace element concentrations to be quantitatively related to bottom water oxygen when the flux of organic carbon is accounted for. Wang et al. (2023) used existing shale trace metal concentration data and machine learning techniques to quantitatively reconstruct oxygen in different Phanerozoic depositional environments (e.g. euxinic basins vs. open ocean OMZs).

### 5.5.2 A better understanding of trace metal delivery to the sediments in the GEOTRACES era

Extensive water column analyses on redox-sensitive trace metals and metal isotopes are essential for revealing their global distribution, source and sink fluxes, and preservation mechanisms in the sediments. The GEOTRACES programme has provided a unique service for mapping dissolved and particulate trace metal (e.g. Mn) and metal isotope distribution in the modern ocean (Schlitzer et al., 2018), allowing for a direct comparison with core-top trace metal and metal isotope measurements. The GEOTRACES data may also advance our understanding of the mass balance and potential isotopic fractionation of multiple trace metals resulting from incorporation or adsorption. These are critical for improving metal isotope mass balance modelling that has been used in quantitative global oxygen reconstructions (e.g. Lau et al., 2019). Future coordinated efforts to expand routine analysis to more redox-sensitive trace metals and metal isotopes (e.g. with robust method development and participation of more research groups), as well as in situ surface sediment collection, would significantly advance proxy development and improve the knowledge of proxy controls and potential caveats.

### 5.5.3 Expanding metal isotope applications in the Cenozoic through proxy development

Apart from the U and Mo isotope systems discussed above, many other “non-traditional” isotope systems are being ac-

tively explored as important redox tracers. Due to the very long residence time of some trace metals compared to seawater (Sect. 6.2.3), other metal isotope proxies have been investigated to study ocean oxygen variability on shorter (e.g. orbital) timescales. For instance, chromium (Cr; residence time of ~10 000 years; Reinhard et al., 2014) isotopes in sediments deposited under sulfidic water columns (e.g. the Cariaco Basin off the Venezuelan coast) may record seawater values due to quantitative Cr removal from water columns (Gueguen et al., 2016; Reinhard et al., 2014). Another promising global oxygen content tracer is provided by thallium (Tl) isotopes, which have been shown to be primarily controlled by the global Mn oxide burial on timescales of < 10 000 years (Nielsen et al., 2011, 2017; Owens et al., 2017). Quantitative Tl removal has been observed in reducing pore waters (with Mn reduction; Ahrens et al., 2021), and a recent core-top calibration suggests that authigenic Tl isotopic compositions can faithfully record the seawater value if pore water is reducing at/near the sediment-water interface, leading to complete Tl sequestration from ambient pore waters (Wang et al., 2022a). As criteria for determining the fidelity of sedimentary Tl isotope records are developed, paleo-reconstructions of seawater Cr, Tl, and other isotopic compositions in the future will ultimately help reveal variations in Tl global ocean content on millennial to orbital timescales, with important implications for marine carbon storage that may have driven the glacial-interglacial transitions (e.g. Wang et al., 2024).

In addition to assessment of global ocean oxygen levels, there is demand for local oxygen reconstructions. This task can be suitably undertaken by proxies with residence times similar to or shorter than the average ocean mixing time of ~1000 years. A promising proxy is cerium (Ce; residence time of the order of ~50–100 years; Alibo and Nozaki, 1999) and Ce isotope ratios. Experiments have shown that oxidative adsorption of dissolved Ce onto Mn oxides can produce ~0.5‰ fractionation in ¹⁴²Ce / ¹⁴⁰Ce, with adsorbed Ce being isotopically light, whereas Ce adsorption onto Fe oxides or Ce oxidation by oxygen produces a smaller Ce isotope fractionation of ~0.2‰ or less (Nakada et al., 2013). This contrasting behaviour in stable Ce isotope fractionation implies a close link between Ce isotope variations and Mn cycling. However, a modern calibration of the Ce isotope system in marine environments is lacking.

## 6 Biomarkers

### 6.1 Organic matter and lipid biomarkers

Organic matter encompasses a wide spectrum of carbon-based compounds of primarily biological origin, which are principally based on carbon-carbon and/or carbon-hydrogen bonds (Killops and Killops, 2005). From a quantitative perspective, most organic matter reaching marine sediments de-
rives from phytoplankton sources from the surface ocean, with additional contributions from bacteria and archaea involved in autotrophic chemosynthesis and heterotrophic processes, particularly in ODZs (Wakeham, 2020). Terrestrially derived organic matter can also be important along continental margins (e.g. Bianchi et al., 2018). Killops and Killops (2005), Eglinton and Repeta (2014), and Peters et al. (2005) provide detailed reviews on the production, composition, degradation, and preservation of organic matter in marine and terrestrial environments.

From a compositional perspective, organic matter largely consists of a few compound classes, including proteins (amino acids), carbohydrates, nucleotides, nucleic acids, and lipids (Killops and Killops, 2005; Peters et al., 2005). Although the former often predominates quantitatively in fresh organic matter (Wakeham et al., 1997), lipids offer by far the largest range of applications in paleoceanography due to their preservation potential in sedimentary systems (Briggs and Summons, 2014; Luo et al., 2019). Lipids include a wide range of compounds that are all characterized by their relatively small molecular size and their mostly hydrophobic nature. This makes them insoluble in water and soluble in organic solvents, such as alkanoic acids; mono-, di-, and triglycerides; waxes; phospho- and glycolipids; lipopolysaccharides; isoprenoids; hopanoids; steroids; terpenes; and pigments, as well as their intact or fragmented fossil remains (Peters et al., 2005). In living cells, lipids play a central role as structural components of membranes for energy storage and as signalling molecules (Hazel and Williams, 1990; van Meer et al., 2008; Harayama and Riezman, 2018). Their intact structure includes a recalcitrant hydrocarbon skeleton that can contain functional moieties such as unsaturations (double bonds) and functional groups (e.g. ester and ether bonds, ketyl, hydroxyl, and carboxyl or amine groups).

The versatility of lipids as the basis of paleoceanographic proxies can be explained by (a) their overall preservation potential in sedimentary systems over geological timescales, (b) their chemotaxonomic and metabolic association with biological sources, (c) their role in controlling cellular physiological processes that lead to lipid remodelling (e.g. degree of unsaturation or cyclization) in response to environmental stressors (e.g. temperature, oxygen, salinity), and (d) the preservation of stable isotope signatures (primarily C and H but also N and S) in their backbone skeletons (e.g. Eglinton and Eglinton, 2008; Eglinton and Repeta, 2014; Peters et al., 2005). Below we provide a brief overview of these processes and how they relate to the reconstruction of redox processes in paleoceanographic studies. Importantly, lipid biomarker applications in paleoceanography follow two approaches: (1) inferring specific source organisms or metabolisms (chemotaxonomy) prevalent in OMZ settings using intentionally biosynthesized compounds and their degradation products and (2) inferring redox conditions using lipid degradation products that only form under oxygen-deficient conditions and may have either ubiquitous sources or no known biolog-

ical sources (orphan biomarkers). In the case of chemotaxonomic approaches, it should be kept in mind that, while other sources of biomarkers may exist or may be discovered in the future (outlined below), independent sedimentological evidence can provide source constraints in a given setting (see Sect. 4).

## 6.2 Lipid preservation and redox potential

The accumulation and preservation of organic material and lipids in marine sediments hinge on a series of physical and biogeochemical factors. These factors include (a) the amount of primary productivity in surface waters, (b) the processes controlling sinking fluxes and attenuation rates of particulate organic matter in its journey through the water column to the seafloor (e.g. availability and composition of ballast material, lateral transport/advection, degree of heterotrophic remineralization), (c) the nature and composition of the organics reaching the sediment-water interface, (d) the redox potential of the depositional environment, (e) the rates of sediment accumulation and burial, (f) the presence of protective minerals (specifically clays), and (g) the availability of reduced sulfur species (e.g. Hedges and Keil, 1995; Blair and Aller, 2012). From their biosynthesis in cells to their preservation in sediments, lipids are subjected to a continuum of post-depositional transformations that modify their physicochemical properties. Initially, these transformations are driven by diagenesis, predominantly microbial enzymatic degradation influenced by the redox potential, which leads to the hydrolysis of polar head groups, and/or the loss of functional groups, and/or the aromatization of ring structures, and the saturation of double bonds (Killops and Killops, 2005; Peters et al., 2005). As sedimentary systems become impacted by tectonic processes and enhanced temperature and pressure gradients, catagenesis and metagenesis lead to changes in the 3D configuration of the molecules (stereochemistry) and, finally, to their thermal cracking (Peters et al., 2005). While the absolute abundance of organic material and lipids decreases along this continuum, the relative abundance of lipids within the total organic material pool increases as a consequence of their higher degradation resistance and preservation potential compared to other compound classes such as carbohydrates and nucleic acids (Briggs and Summons, 2014; Luo et al., 2019). Despite the loss of structural information that lipids endure during degradation, their backbone skeletons preserve diagnostic paleoceanographic information that can be preserved for up to ~1.64 billion years, depending on factors such as oxygen exposure time and thermal maturity (Luo et al., 2019). Thus, since some lipids are more labile (i.e. more prone to degradation) than others (i.e. more recalcitrant), their utility as paleoceanographic proxies is determined by their preservation potential in sedimentary systems over geological timescales (Fig. 6).

The sensitivity of organic matter preservation to bottom water oxygen has long been debated (Pedersen et al., 1992;
Paropkari et al., 1992, 1993). Processes such as oxygen exposure time, the adsorption to mineral phases, and the rate of sediment accumulation have been shown to have the greatest impact (Hedges and Keil, 1995; Hartnett et al., 1998; Burdige, 2007; Zonneveld et al., 2010; Arndt et al., 2013; Hemingway et al., 2019). Organic matter and lipid preservation are enhanced by reducing conditions at the water-sediment interface and within the sediment through (a) reduced exposure time to oxygen-utilizing enzymes, (b) decreased bioturbation, and (c) interactions with reduced sulfur species that lead to lipid sulfurization (e.g. Kohnen et al., 1991). Thus, variable organic matter and lipid preservation, as well as the extent of lipid sulfurization, provide a means of estimating past changes in bottom water oxygen. Empirical studies of organic matter preservation across a range of bottom water oxygen levels in the Arabian Sea have found enhanced preservation of as much as an order of magnitude when bottom water oxygen levels fall below a threshold ranging between 20 and 50 µmol kg⁻¹ (Cowie et al., 2014; Keil and Cowie, 1999; Koho et al., 2013; Rodrigo-Gámiz et al., 2016). Similarly, enhanced accumulation rates and/or preservation of TOC and specific biomarkers under low-oxygen conditions has been found in sediments of the Arabian Sea (Sinninghe Damsté et al., 2002c; Woulds et al., 2009) and off the east coast of the US (Prahl et al., 2001). These studies have shown that the preservation response of biomarkers is nonlinear and that there is a range of sensitivities among different lipid classes to bottom water oxygen.

Anderson et al. (2019) reported that, compared to bulk productivity proxies like opal and $Ba_{xs}$, the accumulation rates of algal lipid biomarkers in sediments deposited during the last glacial period in the central equatorial Pacific Ocean were 5 times greater than during the Early Holocene due to low bottom water oxygen. This interpretation is consistent with independent bottom water oxygen constraints at these sites during the last glacial period (20–50 µmol kg⁻¹) compared to modern bottom water oxygen concentrations (~170 µmol kg⁻¹) (Hoogakker et al., 2018; Umling and Thunell, 2018; Jacobel et al., 2020). Jacobel et al. (2020) demonstrated that when biomarkers are measured in parallel with inorganic proxies for productivity, such as opal and $Ba_{xs}$, it is possible to discriminate between production and preservation as factors causing changes in the concentration or accumulation rate of TOC or of individual compounds, such as oxygen diffusion into the sediments following an increase in bottom water oxygen. The impacts of post-depositional organic matter or biomarker oxidation, a process sometimes referred to as “burndown” (e.g. Colley et al., 1989; Colley and Thomson, 1985; De Lange et al., 2008; De Lange, 1986; Prahl et al., 1989), can be reduced by working at locations with high sediment accumulation rates.

The accumulation and preservation of organic matter and biomarkers can also be enhanced through sulfurization, in which organic matter and organic compounds react with sulfide ($H_2S$, $HS^−$) and/or polysulfides ($S^2_−$), removing func-

tional groups and generating cross-linked polymers that can be relatively resistant to breakdown by microbial exoenzymes (Sinninghe Damsté et al., 1988; Boussafir and Lallier-Vergès, 1997; Van Kaam-Peters et al., 1998). Through these reactions, lipid biomarkers can be bound to high-molecular-weight organic matter (kerogen) via monosulfide (C–S–C) or disulfide (C–S–S–C) bonds (Vairavamurthy et al., 1992; Amrani and Aizenschtat, 2004; Kutuzov et al., 2019). These bonds can be broken during catagenesis (Kelemen et al., 2012) or by chemical desulfurization in the lab (Orr and White, 1990; Prahl et al., 1996; Adam et al., 2000). S-bound and especially disulfide-bound lipids appear to form during very early sedimentation and diagenesis, sometimes prior to the appearance of detectable dissolved sulfide in pore water (Francois, 1987). Early sulfurization can trap biomarker signals before diagenetic reworking and can make these S-bound lipids a relatively high-fidelity archive of biomarker information. In both modern and ancient sediments, S-bound lipid distributions are often distinct from free (extractable) lipids, reflecting important aspects of environmental oxygenation, such as pigments, steroid distributions, and C- or S-isotope compositions (Kohnen et al., 1991; Wakeham et al., 1995; Kok et al., 2000; Rosenberg et al., 2018; Ma et al., 2021). Sulfide in the environment may also contribute to the stabilization of free lipids by reducing double bonds (Hebting et al., 2006). Reconstructions of lipid distributions from sulfidic environments should consider the potential of sulfurization to transform, bias, and/or preserve biomarker information.

Understanding the location and timing of sulfurization also provides insights into the distribution and intensity of anoxia. For example, intervals of enhanced sulfurization and preservation of carbohydrate-derived organic matter in a TOC-rich Jurassic black shale were attributed to photic-zone euxinia during deposition (Boussafir et al., 1995; Van Kaam-Peters et al., 1998; van Dongen et al., 2006). Changes in sulfurization intensity have also been linked to shifts in the distribution of anoxia across OAE2 (Raven et al., 2018). Sulfurization intensity can be approximated by S : C molar ratios, where values greater than about 0.02 exceed the initial sulfur content of most marine photosynthetic biomasses (Francois, 1987). Higher S : C ratios require highly functionalized and therefore relatively young organic precursors prior to sulfurization (Brassell, 1985). This early sulfurization, prior to burial, may impact a relatively large pool of functionalized organic matter (Raven et al., 2018). Subsequently, sulfurization over thousand-year timescales can continue to impact free lipid biomarkers such as tricyclic triterpenoids and steroids (Shawar et al., 2021; Werne et al., 2000; Kok et al., 2000). Different sulfurization products and biogenic organic S can be distinguished through spectroscopy, especially synchrotron X-ray absorption spectroscopy (Kohnen et al., 1989; Amrani and Aizenschtat, 2004; Vairavamurthy, 1998; Raven et al., 2021b), or by using stable isotopes. Organic sulfur in biomass from oxic systems generally has an isotopic compo-
sition similar to that of sulfate in the environment (Kaplan and Rittenburg, 1964; Trust and Fry, 1992), while organic S from sulfurization typically has variable but broadly more ³⁴S-depleted isotopic values (Anderson and Pratt, 1995; Canfield, 2001). In detail, individual organic sulfur compounds have a wide range of stable sulfur isotope (δ³⁴S) values (up to tens of permil) and may preserve additional layers of information about the lipid pool during early diagenesis (Amarani et al., 2005; Raven et al., 2015; Rosenberg et al., 2017; Shawar et al., 2020). Sulfurization indicators can thus complement multi-proxy reconstructions of redox conditions at the same time that they provide insights into taphonomic bias and organic matter burial.

## 6.3 Biomarkers of microbial processes associated with oxygen deficiency

Biomarkers provide chemotaxonomic and metabolic information of source organisms inhabiting water columns and/or sediments impacted by a wide range of redox conditions and electron acceptors. This is particularly important for biomarkers from organisms associated with the cycling of nitrogen (e.g. anammox, ammonia oxidation, nitrogen fixation), sulfur (sulfate reduction, sulfide oxidation), and carbon (methanogenesis and methanotrophy), as well as from those organisms feeding on them. Thus, biomarkers provide qualitative redox information that ranges from fully oxygenated conditions to oxygen-deficient, anoxic non-sulfidic, and anoxic sulfidic/euxinic conditions (both within and below the photic zone). We refer to Table S1 for a tabular summary of biomarkers described in Sect. 6.3.1 to 6.3.2.

### 6.3.1 Biomarkers for nitrogen cycling in ODZs

Oxygen-deficient environments are hotspots for microbial processes involved in the removal of fixed nitrogen, such as denitrification and other dissimilatory nitrogen transformations, anaerobic ammonium oxidation (anammox) coupled to nitrite reduction, and nitrite-dependent anaerobic methane oxidation (n-damo) (Lam and Kuypers, 2011; Thamdrup, 2012). Accordingly, the presence of lipids synthesized by bacteria with these anaerobic metabolisms can be used to infer hypoxic or anoxic conditions in the past. For detailed reviews on the use of biomarkers for nitrogen cycling, we recommend the recent work by Rush and Sinninghe Damsté (2017) and Kusch and Rush (2022).

#### Anammox

Anammox, the anaerobic oxidation of $NH_4^+$ to $N_2$ using $NO_2^-$, is the only nitrogen loss process for which several specific biomarkers have been identified. Ladderanes are highly unusual lipids with moieties consisting of cyclobutane rings (Sinninghe Damsté et al., 2002a). They make up the cell membrane of the anammoxosome, the specialized organelle in which the anammox process takes place, and substantially

reduce proton permeability (Moss et al., 2018). A suite of ladderane fatty acids and their short-chain oxic degradation products are available to trace anammox. However, ladderanes do not preserve well (Rush et al., 2011), and, thus far, the oldest sedimentary record containing ladderanes extends only 140 kyr (Jaeschke et al., 2009). More recently, a unique bacteriohopanepolyol (BHP) biomarker for marine anammox *Candidatus Scalindua* profunda was identified (Schwartz-Narbonne et al., 2020): a bacteriohopanetetrol (BHT) isomer with unknown stereochemistry, BHT-x (Rush et al., 2014; Schwartz-Narbonne et al., 2020), has strongly depleted δ¹³C values in sediments consistent with fractionation associated with the reductive acetyl-CoA pathway used by anammox (Hemingway et al., 2018; Lengger et al., 2019) and has a distinct niche in the nitrite maximum of stratified water column settings (e.g. Kusch et al., 2022; Matys et al., 2017). In comparison to ladderanes, BHPs preserve well, and BHT has been detected in samples as old as ca. 50–55 million years (van Dongen et al., 2006; Talbot et al., 2016). In older rocks, BHPs may survive as hydrocarbons after the loss of hydroxyl functionalities due to reducing processes or after decarboxylation reactions that also shorten the hopanoid side chain. The diagnostic value of these resulting hopenes and hopanes is reduced in comparison to the original lipid. So far, BHT-x has been successfully used to trace OMZ conditions in (non-dated) sediment records in the Benguela Upwelling System (van Kemenade et al., 2022) during the last glacial in the Gulf of Alaska (Zindorf et al., 2020) and during Pliocene/Quaternary sapropel formation in the Mediterranean (Rush et al., 2019; Elling et al., 2021). For paleo-oceanographic purposes, sedimentary ladderanes and BHT-x should primarily capture the water column anammox signal (nitrite maximum), which can be orders of magnitudes higher than the benthic background signal (Rush et al., 2012).

#### Nitrite-dependent anaerobic methane oxidation (n-damo)

More recently, bacteria performing anaerobic methanotrophy have been detected. *Candidatus Methylomirabilis oxyfera* is an exceptional methanotroph that produces its own oxygen via the production of NO by $NO_2^-$ reduction (Ettwig et al., 2010), also known as *n*-damo. Biomarkers of *n*-damo include bacteriohopanehexol, 3Me-bacteriohopanehexol, and 3Me-bacteriohopanepentol (Kool et al., 2014) as well as the novel demethylated hopanoids 22,29,30-trisnorhopan-21-ol, 3Me-22,29,30-trisnorhopan-21-ol, and 3Me-22,29,30-trisnorhopan-21-one (Smit et al., 2019). Although not common, bacteriohopanehexol is also produced by thermophilic *Alicyclobacillus acidoterrestris* (Rezanka et al., 2011) and the bacterial symbiont of a marine *Petrosia* sponge (Shatz et al., 2000). Thus, the C-3 homologues of bacteriohopanehexol and bacteriohopanepentol as well as the trisnorhopanes may be better indicators of the presence of *Candidatus Methylomirabilis oxyfera* and *n*-damo. Although there seem to be
several *n*-damo biomarkers, it should be noted that the role of *n*-damo bacteria (NC10 phylum) in marine ODZs is still not well constrained (e.g. Padilla et al., 2016). However, to date the presence of the above-mentioned biomarkers in marine sediments can likely be interpreted to indicate anoxia.

### 6.3.2 Feedback mechanisms to nitrogen loss

Removal of nitrogen in OMZ settings causes imbalances in N:P ratios that can promote/intensify aerobic processes involved in the nitrogen cycle. For instance, in the geo-logic record, enhanced diazotrophy, the primary source of bioavailable nitrogen in the ocean (Hutchins et al., 2009), has been invoked to have sustained biological productivity dur-ing times of intensified ocean deoxygenation and consequent fixed nitrogen loss (e.g. Kuypers et al., 2004). Accordingly, biomarker evidence of important feedback mechanisms to ni-trogen loss (e.g. diazotrophy and nitrification) can provide context for paleoceanographic data and shed light on past OMZ-related biogeochemical cycles.

Biomarkers for reconstructing nitrogen fixation have long been sought after. To date, there are no known biomarkers commonly produced by cyanobacteria or the subset of cyanobacteria that can fix nitrogen and are the major dia-zotrophs in the ocean (e.g. Talbot et al., 2008; Sáenz et al., 2012). Therefore, molecular evidence of N₂ fixation and den-itrification has mostly been based on the characteristic ki-netic isotope fractionation effects (e.g. Sigman et al., 2009b) preserved in the δ¹⁵N values of nitrogen-containing organics such as pigments and proteins (e.g. Sachs and Repeta, 1999; Ohkouchi et al., 2006; Higgins et al., 2010, 2012; Junium et al., 2015). For example, the consistent observation of low δ¹⁵N values of chlorophyll-derived tetrapyrroles in ancient sediments deposited under widespread anoxic and euxinic conditions (e.g. OAEs) and the abundance of chlorophyll-derived degradation products suggest a direct link between surface water N₂ fixation and water column N loss processes (e.g. Junium and Arthur, 2007; Ohkouchi et al., 2006). Re-cent studies also suggest that δ¹⁵N values of amino acids (δ¹⁵N<sub>AA</sub>) may be useful tools to study water column nitro-gen dynamics (McCarthy et al., 2013; Batista et al., 2014) since phenylalanine δ¹⁵N values show a good relationship with established δ¹⁵N proxies such as bulk sediment and foraminifera-bound N (Li et al., 2019) (see also Sect. 7; nitrogen isotopes). However, for a subset of diazotrophic cyanobacteria, so-called heterocystous cyanobacteria, spe-cific heterocyst glycolipids (HGs) with C₅ and C₆ sugar head groups have now been identified (Bauersachs et al., 2009a, b; Bauersachs et al., 2010). HGs comprise the inner-most laminated layer of the heterocysts, forming a protec-tive envelope for the oxygen-sensitive nitrogenase enzyme (e.g. Gambacorta et al., 1998). C₅ sugar HGs are proposed to be specific biomarkers for marine endosymbiotic heterocys-tous cyanobacteria, while C₆ sugar HGs occur in free-living heterocystous cyanobacteria (Schouten et al., 2013b; Bale et

al., 2015). Enhanced deposition of C5 and C6 HGs in the Mediterranean Sea during Plio-Pleistocene sapropel events has been linked to anoxia, indicating that diazotrophy by het-erocystous cyanobacteria was an important feedback to ni-trogen loss (Bale et al., 2019; Elling et al., 2021). Hetero-cystous cyanobacteria occur mostly in brackish water bodies (e.g. Baltic Sea) but are rare in the open ocean (except as symbionts of diatoms; Stal, 2009; Zehr, 2011). However, di-zotrophy is also observed in open ocean ODZ systems that are associated with enhanced upwelling and primary produc-tion, such as the eastern tropical Pacific (White et al., 2013; Loescher et al., 2014; Jayakumar et al., 2017). It is unknown whether HGs can also track heterocystous cyanobacteria in these environments in the past or present. Thus, tetrapyrrole δ¹⁵N values may still provide the most unequivocal evidence of N₂ fixation in the past. Since they preserve well over long timescales (the oldest tetrapyrroles date back to 1.1 billion years ago; Gueneli et al., 2018), the nitrogen isotopic com-position of these molecules or their smaller maleimide frag-ments (e.g. Grice et al., 1996) can be used to gauge N₂ fixa-tion over much of Earth's history.

Nitrification, the aerobic transformation of ammonia (NH₄⁺) to nitrite (NO₃⁻), is performed either as a two-step process by ammonia-oxidizing bacteria (AOB), ammonia-oxidizing archaea (AOA), and nitrite-oxidizing bacteria (NOB) or as a one-step process by ammonia-oxidizing bacteria (comammox). Although nitrification is considered an obligately aerobic process, AOA and NOB persist in suboxic and anoxic waters, and two novel *Nitrospina*-like lineages (NOB) have been found and implicated in nitrite oxidation in ODZs (Sun et al., 2019). Thus, for the majority of known AOA and NOB, active nitrification under low-oxygen condi-tions requires a source of cryptic oxygen, i.e. the presence of short-lived oxygen-bearing intermediates that typically oc-cur below detection limits but provide important substrates (Kappler and Bryce, 2017). For AOA, internal oxygen pro-duction has been observed as a response to anoxia (Kraft et al., 2022). AOA are the dominant sources of glycerol dialkyl glycerol tetraethers (GDGTs) to marine sediments and pro-duce the specific biomarker crenarchaeol (Sinninghe Damsté et al., 2002b), methoxy archaeol (Elling et al., 2014, 2017), and specific quinones (MK<sub>6:0</sub> and MK<sub>6:1</sub>; Elling et al., 2016). Crenarchaeol has been shown to track Thaumarchaeota in the suboxic zones of modern (e.g. Wakeham et al., 2007; Sol-lai et al., 2015; Kusch et al., 2021) and paleo systems, par-ticularly during times of ocean deoxygenation, such as dur-ing Mediterranean sapropel deposition (Menzel et al., 2006; Polik et al., 2018). As such, increased deposition of cre-narchaeol relative to organic matter may be useful for trac-ing intensified suboxic-anoxic nitrogen cycling (Rush and Sinninghe Damsté, 2017; Elling et al., 2020). However, the presence of crenarchaeol alone cannot be used to infer sub-oxic conditions. AOB, NOB, and comammox do not seem to synthesize chemotaxonomically specific lipids. Known lipids of AOB include generic BHPs and unsaturated fatty acids
(Sakata et al., 2008), and some hopanoids produced by NOB
have not previously been found in other bacteria (Rush and
Sinninghe Damsté, 2017; Elling et al., 2022).

6.4 Biomarkers for sulfur cycling in ODZs

In addition to the abiotic sulfurization mechanisms described
above, ODZs are characterized by active sulfide oxidation
and sulfate reduction mediated by diverse bacteria (e.g. Call-
beck et al., 2021; van Vliet et al., 2021). These sulfur
metabolisms are present not only in sulfide-rich anoxic sed-
iments and euxinic water columns, but also in open ocean
ODZs. Particle microniches also harbour a cryptic sulfur
cycle (Canfield et al., 2010; Raven et al., 2021a). Diverse
biomarkers with high preservation potential have been identi-
fied for various sulfide-oxidizing bacteria (SOB) and sulfate-
reducing bacteria (SRB) and allow for detailed reconstruc-
tions of water column stratification in paleoceanographic
studies.

6.4.1 Sulfide oxidation

Phototrophic green sulfur bacteria (GSB; Chlorobiaceae) and purple sulfur bacteria (PSB; Chromatiaceae) are the principal microbes metabolizing reduced sulfur species, such as H₂S, whilst fixing carbon through anoxygenic photosynthesis. As such, they occupy anoxic environmental niches with access to light and H₂S, amongst other sulfur species (Summons and Powell, 1987). Apart from benthic microbial mats at shallow water depths, this involves euxinic photic zones of marine water columns. Characteristically, green and purple sulfur bacteria biosynthesize diaromatic carotenoids that function as accessory pigments. Isorenieratene and chlorobactene, as well as their fossilized equivalents isorenieratane and chlorobactane, are commonly used as indicators of the presence and activity of green- (chlorobactane) and brown-pigmented (isorenieratane) species of Chlorobiaceae during sediment deposition, providing clues about the depth of the chemocline given that brown-pigmented Chlorobiaceae are adapted to lower irradiance than their green-pigmented relatives (Summons and Powell, 1987; Schaeffer et al., 1997; French et al., 2015). In very shallow chemoclines, the relative abundance of anoxygenic phototrophy is typically skewed towards a higher proportion of PSB, which characteristically biosynthesize the monoaromatic carotenoid okenone that can survive in sediments as the fossil-equivalent okenane (Brocks and Schaeffer, 2008). All of the saturated C₄₀ carotenoids can survive for exceedingly long time spans and have been detected in sediments up to 1.64 Ga in age (see review and updated analytical method in French et al., 2015). Once subjected to thermal breakdown during sedimentary burial, the methylation pattern of the remaining aryl isoprenoids (2,3,4 vs. 2,3,6 substitution pattern) can still yield clues to the biological precursor, whereas the relative abundance of longer versus shorter aryl isoprenoid chains may allow for distin-

guishing long-lived and persistent euxinia from short-lived
and episodic photic-zone euxinia (Schwark and Frimmel,
2004). Using the reverse tricarboxylic acid cycle during car-
bon fixation, the biomass of GSB may also be recognized
by its characteristic enrichment in ¹³C compared to that of
oxygenic phototrophs, whilst its bacteriochlorophyll c/d/e
pigments can be recognized intact and after breakdown
to maleimides, such as 3-isobutyl-4-methylmaleimide (e.g.
Grice et al., 1996; Naeher et al., 2013).

6.4.2 Sulfate reduction

Sulfate reduction is a heterotrophic anaerobic bacterial path-
way leading to the formation of H2S, which can fuel the
cryptic sulfur cycle in offshore ODZs by supplying reac-
tive sulfur species as intermediates for other redox reactions
(Callbeck et al., 2021). A group of compounds commonly
associated with SRB are non-isoprenoid 1-O-monoalkyl or
2-O-monoalkyl glycerol ethers (MAGEs) and 1,2-O-dialkyl
glycerol ethers (DAGEs). They have been identified in hyper-
thermophilic bacteria and commonly occur in settings influ-
enced by hydro/geothermal activity (e.g. Bradley et al., 2009)
or seep systems hosting consortia of anaerobic methane-
oxidizing archaea (ANME) and SRB (Niemann and Elvert,
2008). Hernandez-Sanchez et al. (2014) also identified 1-O-
MAGEs in suspended particulate matter sampled from oxy-
genated surface waters and suggested a role of bacteria other
than SRB in the production of these lipids. However, re-
cent evidence of sulfate reduction in sinking marine parti-
cles (Raven et al., 2021a) can explain the observation of SRB
or 1-O-MAGEs in oxygenated surface waters and sediments
(e.g. Hernandez-Sanchez et al., 2014; Teske et al., 1996).
SRB belonging to the *Desulfosarcina/Desulfococcus* group
(syntrophic partners of the ANME-1 and ANME-2 clades,
except ANME-2d) and *Desulfobulbus* spp. (syntrophic part-
ner of the ANME-3 clade) also produce characteristic alka-
noic acid fingerprints with strong ¹³C depletion, including
C₁₆:₁ω₅ (C₁₆:₁Δ¹²), cy-C₁₇:₀ω₅,₆, and C₁₇:₁ω₆ (Niemann
and Elvert, 2008). No biomarkers are known for sulfur-
disproportionating bacteria, which perform reverse sulfate
reduction.

6.5 Biomarkers for carbon cycling in ODZs

Oxygen-deficient conditions in the ocean are also inti-
mately linked to the methane cycle via both the genera-
tion (methanogenesis) and the utilization (methanotrophy)
of methane. Methanogenesis and anaerobic methanotrophy
are performed by anaerobic archaea using sulfate, nitrate,
Fe, and Mn as electron acceptors (for a recent overview,
see Guerrero-Cruz et al., 2021). Methane is also respired by
aerobic methane-oxidizing bacteria (MOB). Although these
MOB are aerobes, their lipids are useful proxies in paleo-
ceanography since MOB are typically present at oxic-anoxic
transitions where both methane and oxygen are available
and can thrive under oxygen-deficient conditions (Guerrero-Cruz et al., 2021). In addition, the utilization of methane is recorded by strongly ¹³C-depleted biomarker signatures irrespective of their chemotaxonomic specificity (e.g. Jahnke et al., 1999).

### 6.5.1 Methanogenesis

Methane in the ocean is primarily produced in anoxic marine sediments, although aerobic sources also exist (Metcalf et al., 2012; Bižić et al., 2020). In sediments, methanogenesis is performed by strictly anaerobic and primarily hydrogenotrophic and acetoclastic Euryarchaeota (Ferry and Lessner, 2008). Methanogenic Euryarchaeota primarily produce isoprenoid tetraethers without cyclic moieties (GDGT-0) and the isoprenoid diether archaeol (Koga et al., 1993, 1998). Ratios > 2 of GDGT-0 over the thaumarchaeal isoprenoid tetraether crenarchaeol have been proposed as a proxy for methanogens in the paleo record (Blaga et al., 2009), although these ratios have to be interpreted in the context of the biomarker assemblage due to the presence of GDGT-0 in many non-methanogenic archaea (Schouten et al., 2013a; Elling et al., 2017).

Likewise, *Methanothermococcus thermolithotrophicus* has been shown to synthesize hydroxylated GDGTs (OH-GDGT) with 0–2 pentacyclic moieties (Liu et al., 2012). Although the lipids mentioned above are not exclusive to methanogens, other sources such as methanotrophic Euryarchaeota or Thaumarchaeota typically have much more diverse GDGT fingerprints (Schouten et al., 2013a; Elling et al. 2017) and also synthesize additional lipids (see below). Since methanogenesis is performed in situ, for paleoceanographic purposes it is crucial to identify in situ overprints. One approach to distinguish between paleo and in situ signals is the screening for biomarkers that imply metabolic activity, such as the functional quinone analogues methanophenazines (Abken et al., 1998; Elling et al., 2016), which (to date) have only been shown to be produced by the order Methanosarcinales or coenzyme F430, the cofactor of the methyl coenzyme M reductase possessed by all methanogens (Kaneko et al., 2021). However, it should be noted that anaerobic methanotrophic archaea are also suspected of producing these biomarkers (although direct evidence is still missing).

### 6.5.2 Methanotrophy

Anaerobic methanotrophy with sulfate (reverse methanogenesis) is performed by ANME archaea in syntrophic consortia with sulfate-reducing bacteria (Boetius et al., 2000). ANME biomarkers include isoprenoids such as tetramethylhexadecane (crocetane; ANME-2), pentamethylicosane (PMI), and unsaturated PMIs, as well as archaeol, *sn*2-hydroxyarchaeol, and *sn*3-hydroxyarchaeol (Koga et al., 1993, 1998; Elvert et al., 1999; Thiel et al., 1999; Hinrichs et al., 2000,

2003). Furthermore, ANME ecotypes (classified as ANME-1, ANME-2, and ANME-3) seem to be discernible by *sn*2-hydroxyarchaeol/archaeol ratios and the δ¹³C signature of archaeol, the alkanoic acid signature of the SRB partners (Blumenberg et al., 2004; Niemann and Elvert, 2008), and specific intact polar lipid (IPL) compositions (e.g. Rossel et al., 2011), although the IPL characteristics might be lost in the paleo record. However, the remaining core lipid GDGT signature may aid the identification of ANME in the paleo record. GDGT-2/crenarchaeol ratios exceeding the threshold of 0.2 could indicate the presence of ANME (Weijers et al., 2011). The characteristic depletion in ¹³C of ANME biomarkers can be traced in the paleo record.

Methane produced under anoxic conditions can also be utilized by aerobic MOB, and its presence in paleoceanographic records indicates a methane-rich environment. Aerobic MOB synthesize a range of characteristic lipids, including a suite of amino-functionalized BHPs and their respective unsaturated and C-3 methylated homologues. Aminopentol is considered a characteristic biomarker for Type I gammaproteobacterial MOB, and aminotetrol is commonly produced by Type II alphaproteobacterial MOB (e.g. Rohmer et al., 1984; Jahnke et al., 1999; Talbot et al., 2001). They also produce structurally similar methyl carbamate (MC) BHPs (MC-pentol and MC-tetrol), which seem to be much more common in methane-influenced marine environments that often lack aminopentol (Rush et al., 2016). It should be noted that minor amounts of aminopentol and aminotetrol are also produced by SRB of the *Desulfovibrio* genus (Blumenberg et al., 2006, 2012), NOB (Elling et al., 2022) and several terrestrial thermophilic bacterial species (Kolouchová et al., 2021). Likewise, small amounts of MC-triol have recently been identified in cultures of *Nitrobacter vulgaris* and marine *Nitrococcus mobilis* (Elling et al., 2022). C-3 methylation of hopanoids alone can no longer unequivocally be linked to AOM unless confirmed by depleted δ¹³C values. The gene for C-3 methylation is not present in all methanotrophs but is present in various non-methanotrophic bacteria (Welander and Summons, 2012); C-3 methylated BHPs accumulate in the euxinic Black Sea water column (Kusch et al., 2022), and other known sources include the phototrophic purple nonsulfur bacterium *Rhodopila globiformis* (Mayer et al., 2021). However, the presence of MOB may also be confirmed by other biomarkers such as alkanoic acids and quinones. Type I MOB produce methylene ubiquinone MQ₈:7 (Nowicka and Kruk, 2010), which has been observed in the suboxic and anoxic zones of the Black Sea (Becker et al., 2018), and Methyllococcaceae synthesize C₁₆:₁Δ⁸c (C₁₆:₁ω⁸c) and C₁₆:₁Δ¹²t (C₁₆:₁ω⁵t) alkanoic acids (e.g. Bodelier et al., 2009). Type II MOB of the Methylocystaceae and Beijerinckiaceae are characterized by C₁₈:₁Δ¹⁰c (C₁₈:₁ω⁸c) (e.g. Bodelier et al., 2009).
## 6.6 Non-specific/orphan biomarkers from oxygen-deficient depositional settings

### 6.6.1 Redox-controlled processes

Independent of their specific biological source, various lipids will undergo diagenetic molecular modifications that are principally controlled by environmental redox chemistry. One of the first established indicators of oxic versus anoxic conditions was the ratio of pristane (Pr) to phytane (Ph) – C₁₉ and C₂₀ isoprenoid hydrocarbons that derive from the phytol side chain of chlorophyll by oxidative decarboxylation or by reduction, respectively (Rontani and Volkman, 2003). Strongly elevated values (e.g. > 4) are observed under oxic conditions, whereas values ≪ 1 are found under anoxic conditions, yet the use of the Pr/Ph index is complicated by alternative (non-chlorophyll-derived) sources of both Pr and Ph (e.g. Goossens et al., 1984), as well as by questions of organic matter transport pathways and oxygen exposure (e.g. Ten Haven et al., 1987). A hopanoid-based indicator involving the relative abundance of long-chain C₃₁–C₃₅ homohopanes over C₃₀ hopanes (known as the homohopane index) follows a similar rationale and assumes that longer side chains are preferentially preserved in sediments under reducing conditions (Peters et al., 2005). Similarly, phototroph-derived chlorophyll is commonly used as an indicator of primary productivity (Carpenter et al., 1986; Harris et al., 1996). However, certain degradation products, such as pyropheophytin and steryl chlorin esters, are only formed under anoxic conditions (Szymczak-Żyla et al., 2008), and the proportional abundance of these chlorophyll degradation products has been proposed as a proxy for bottom water anoxia (e.g. Szymczak-Żyla et al., 2017).

### 6.6.2 Orphan biomarkers

A range of lipids have been shown to be associated with OMZ settings for which the source organisms are unknown. Although these lipids cannot be linked to specific taxa or metabolisms, these orphan biomarkers can still be useful indicators for paleoceanographic purposes. One common orphan biomarker in sediments from anoxic settings is the isoprenoid 19,23,27,31-octamethyldotriacontane (lycopane), for which methanogenic archaeal (Brassell et al., 1981) and phototrophic (Wakeham et al., 1993) origins have been suggested. For paleoceanographic purposes, the lycopane / C₃₁ n-alkane ratio has been suggested as a proxy for paleoxygen levels (Sinninghe Damsté et al., 2003), although it must be applied with caution in areas where plant wax input is large.

Derivatives of branched GDGTs, such as overly branched GDGTs (OB-GDGTs), are produced in specific patterns in anoxic marine zones (Liu et al., 2014; Xie et al., 2014) and have been interpreted as biomarkers for OMZ presence in the paleo record (Connock et al., 2022). Yet, the sources of these

lipids, oxygen thresholds for their production, and potential production in sediments remain unstudied.

## 6.7 Analyses and resources required

For biomarker analyses, sediments are extracted with organic solvents to recover the total lipid extract (TLE; or bitumen), which is further processed to separate compound classes that are subsequently analysed using mass spectrometry techniques. For extraction, pure organic solvents or mixtures that contain water and/or buffers are used with ultrasonication, Soxhlet, or automated pressurized systems such as ASE and microwave. Extraction is typically followed by wet-chemical processing to obtain polarity fractions, and it may be necessary to remove elemental sulfur to avoid interference during chromatography or release S-bound organics that would otherwise evade detection (Kohnen et al., 1991).

Small non-polar compounds can be analysed using gas chromatography–mass spectrometry (GC-MS), whereas larger and/or more polar compounds require the use of liquid chromatography–mass spectrometry (LC-MS) techniques, unless they are derivatized and/or cleaved from polar functional groups. GC isotope ratio mass spectrometry (GC-IRMS) systems are used to obtain compound-specific stable isotope values (δ¹⁵N, δ¹³C, δ²H), or compounds are isolated using LC and subsequently analysed using spooling-wire micro-combustion IRMS (Pearson et al., 2016) or nano-elemental analyser (EA)-IRMS analysis (e.g. Kusch et al., 2010; Ogawa et al., 2010).

## 6.8 Major analytical advancements

Major advancements in the field of organic geochemistry have been made since the introduction of high-resolution accurate mass–mass spectrometry (HRAM-MS) techniques, including quadrupole time of flight (Q-ToF), Orbitrap, and Fourier transform ion cyclotron resonance (FT-ICR). Paired with ultra-high-performance LC systems, improved column materials and chemistries (e.g. core shell, HILIC), Q-ToF and Orbitrap platforms have opened up a new window into the lipidomics of environmental samples (e.g. Hopmans et al., 2021; Wörmer et al., 2013). Orbitrap technology specifically offers substantial further analytical potential that includes compound-specific isotope analysis and position-specific isotope analysis (Eiler et al., 2017). Recent analytical advancements have also been made using scanning techniques. Matrix-assisted laser desorption/ionization (MALDI)-FT-ICR-MS allows for mapping of spatial biomarker abundances in situ, which facilitates the obtainment of ultra-high-resolution records from sediment cores (e.g. Alfken et al., 2021). For a comprehensive overview of high-resolution analytical organic geochemical methods, we refer to Steen et al. (2020).
## 6.9 Future prospects

Biomarkers are excellent tools in paleoceanography due to their preservation potential in sediment and their biological and environmental associations. Open ocean ODZs in general are high-productivity systems that foster preservation of high amounts of organic matter. Moreover, the high preservation potential of lipids leads not only to high abundances of biomarkers in the paleo record, but also to a diverse pool of compounds. This structural diversity allows for the use of comprehensive biomarker assemblage approaches to reconstruct the water column structure (e.g. Connock et al., 2022; Dummann et al., 2021). Biomarkers also preserve well under conditions that lead to the absence of calcareous microfossils or trace metals in sediments, such as carbonate undersaturation or oxygen exposure. Nonetheless, challenges and new frontiers remain for the biomarker community. Most profoundly, biomarker proxies have yet to allow for the quantitative reconstruction of oxygen concentrations in the past. Indirectly, the half-maximal inhibitory concentration (IC₅₀) of oxygen in specific organisms provides a means to infer maximum oxygen concentrations. For example, anammox bacteria have an upper oxygen limit of ~20 µmol kg⁻¹ (Kalvelage et al., 2011), and a recent study suggests that BHT-x ratios (the normalized abundance of BHT-x / [BHT + BHT-x]) of ≥ 0.2 indicate oxygen concentrations < 50 µmol kg⁻¹ (van Kemenade et al., 2022). Recently, Kim and Zhang (2023) demonstrated how the methane index can be used to calculate sedimentary methane fluxes. Analogous proxies to reconstruct oxygen are lacking, as a range of lipids are produced in the water column and in sediments (e.g. anammox, sulfate reduction), and different biomarkers may sometimes provide divergent paleoenvironmental information (e.g. in settings receiving input from land or affected by lateral input). In addition to making the available biomarker tools more quantitative, efforts of the biomarker community are also focused on exploiting the potential afforded by technological advances to expand the utility of biomarker applications in paleoceanography. New isotope tools, such as compound-specific ¹⁵N analysis of amino acids, as well as the discovery of new biomarkers through lipidomic approaches in combination with (meta)genomics, offer large potential to trace oxygen limitation in more detail. These new tools may also aid the identification of the sources of orphan biomarkers that accumulate under anoxic conditions. Identifying the source organisms will improve our understanding of the metabolism or ecological niche recorded by these lipids and, ultimately, the paleoenvironmental conclusions drawn from them.

<figure><img src="image_8.png" /><figcaption>**Figure 6.** (a) Idealized water column and sediment redox zone tion and associated metabolisms and their biomarker signatures discussed in the text. The dashed lines indicate species and metabolisms that may be present in specific settings. (b) Biomarker preservation timescale. The solid lines indicate timescales on which distinct compound classes are preserved. Note that observations are less frequent where the scale extends to the oldest known record.</figcaption></figure>

# 7 Nitrogen isotopes

## 7.1 Introduction

Nitrogen (N) has two stable isotopes, ¹⁴N and ¹⁵N. The lighter isotope, ¹⁴N, comprises 99.63 ± 0.02% of the N found on Earth. Natural processes in the ocean discriminate between the two isotopes, leading to subtle changes in the ¹⁵N / ¹⁴N ratio of different nitrogen compounds. This is reported in delta notation, with atmospheric N₂ as the reference
material:

$$ \delta^{15}\text{N}(\text{‰ vs. Air}) = \left( \frac{(^{15}\text{N}/^{14}\text{N})\text{Sample}}{(^{15}\text{N}/^{14}\text{N})\text{Air} - 1} \right) \cdot 1000 $$

The most widely used technique for N isotopic analysis relies on online combustion of samples to N₂ using an elemental analyser (EA) coupled with isotope ratio mass spectrometry (IRMS). This method requires a typical sample size of 1–2 mmol N per analysis, with analytical precision near 0.1% (1 standard deviation), and has mostly been applied to bulk sedimentary N (Robinson et al., 2012; Tesdal et al., 2013). Ongoing innovations, for example through cryo-focusing of the resulting N₂ sample gas, allow for greatly reduced sample sizes (~25–40 nmol N) (Polissar et al., 2009). Specific organic compounds, amino acids, and other polar compounds typically require a derivatization process for gas-chromatograph-based analysis. Amino acids can also be isolated using liquid chromatography followed by micro-combustion to N₂, which is commonly applied prior to the IRMS analysis of the resulting N₂ (Ohkouchi et al., 2017; Ishikawa et al., 2022).

In addition to N₂, N₂O can be analysed by IRMS for N isotopic analysis. As N₂O has a much lower atmospheric background relative to N₂, the sample size requirement can be further reduced when N₂O is the final analyte. Methods have been developed for the conversion of nitrate (NO₃⁻) and nitrite (NO₂⁻) to N₂O, preceded by chemical procedures to convert different N forms to NO₃⁻/NO₂⁻ (Sigman et al., 2001; McIlvin and Altabet, 2005; Weigand et al., 2016). The N₂O-based method, using denitrifying bacteria, has a typical sample size requirement of 2–5 nmol N, with analytical precision better than 0.2‰ (Weigand et al., 2016). This technique has expanded the range of sample types accessible to isotopic analyses and has successfully been applied to a range of fossil-bound N, including diatoms, foraminifera, and corals, among others.

Nitrogen cycling is closely linked to the dissolved oxygen content in the ocean. In ODZs, where water column oxygen concentrations are lower than roughly 5 µmol kg⁻¹, the bacterial reduction in NO₃⁻ to N₂, also known as denitrification, removes NO₃⁻ from the ocean (Sigman et al., 2009b). Denitrification strongly discriminates against the heavier isotope, progressively increasing the δ¹⁵N of the remaining NO₃⁻ pool. Culture and field studies show an isotope effect between 15‰ and 25‰ for denitrification (Sigman and Fripiat, 2019). Recent work suggests that anaerobic ammonium oxidation (anammox), where NO₂⁻ is used to oxidize ammonium (NH₄⁺) to N₂, could also lead to an increase in the N isotope composition of NO₃⁻ (δ¹⁵Nnitrate) from suboxic environments, where extensive NO₂⁻ oxidation co-occurs with anammox (Brunner et al., 2013; Casciotti, 2016). Multiple factors influencing the net effect of anammox on δ¹⁵N are

still under investigation (Kobayashi et al., 2019). However, since denitrification and anammox typically coexist in suboxic environments, isotope effects estimated from previous field studies should include both processes. For the purpose of this review, we only refer to denitrification but note that denitrification likely coexists with anammox. Because of the low-oxygen threshold for denitrification and anammox, the two processes only occur in ODZs, which represent the most oxygen-depleted areas within the broader low-oxygen zones (OMZs).

In contrast to water column denitrification, sedimentary denitrification leads to little increase in the δ¹⁵Nnitrate in the overlaying water column because the isotopic discrimination is minimized by nearly complete consumption of NO₃⁻ at the site of denitrification within sediment pore waters, leading to a small overall isotope effect (less than 3‰) (Sigman and Fripiat, 2019). In highly productive ocean margin environments, such as sub-Arctic and Arctic shelves, sedimentary denitrification has been reported to have a somewhat greater effect (up to 8‰) (Granger et al., 2004), although the effect on water column NO₃⁻ remains small.

In summary, denitrification/anammox leaves a strong imprint on the isotopic composition of the residual NO₃⁻ in the water column of ODZs (Cline and Kaplan, 1975; Liu and Kaplan, 1989). In those regions, isotopically enriched NO₃⁻ can be upwelled and taken up by primary producers. The organic N produced in the surface ocean is exported, part of which is remineralized below the euphotic zone, spreading the high δ¹⁵N signal of denitrification/anammox further away from the ODZs and suboxic environments through lateral transport (Sigman et al., 2009a, b). The organic N that is not respired can carry the resulting isotopic enrichment into the underlying sediment and, if preserved well during sinking and burial, can be used as a recorder for changes in δ¹⁵Nnitrate and thus changes in the extent of ODZs in the past (Altabet et al., 1995; Ganeshram et al., 1995).

The δ¹⁵N of sedimentary archives is a semi-quantitative proxy for oxygen content in the water column. Several challenges must be overcome to quantitatively calibrate the proxy. This will require understanding of (i) uncertainties in the denitrification/anammox isotope effect, (ii) controls on denitrification rates, and (iii) influences of other processes such as NO₃⁻ consumption and N₂ fixation on δ¹⁵Nnitrate. First, the current estimate of the isotope effect associated with denitrification/anammox has a large range. Earlier field-based estimates for the isotope effect often led to overestimation by comparing NO₃⁻ from dysoxic environments with deep-ocean NO₃⁻ (Marconi et al., 2017). This can be improved with new analyses with the assistance of modelling and culture studies. Second, controls on the overall rate of denitrification in dysoxic environments are not fully constrained, although it is most likely tied to organic carbon supply (Ward et al., 2008). Field and modelling studies are needed to understand the relationship between the rate of denitrification/anammox, the size of the ODZs, and the oxy-
gen content in the OMZ. Third, other biological processes occurring in the upper water column, especially $NO_3^-$ consumption and $N_2$ fixation, can affect $\delta^{15}N_{nitrate}$ when it is upwelled from dysoxic waters to the surface ocean (e.g. Farrell et al., 1995). The three major modern OMZs in the eastern tropical North Pacific and South Pacific and the Arabian Sea are all currently characterized by strong upwelling; high productivity; and, in many cases, incomplete $NO_3^-$ utilization in the surface ocean. Because phytoplankton preferentially assimilate $^{14}N$ relative to $^{15}N$, incomplete utilization of surface $NO_3^-$ would elevate surface $\delta^{15}N_{nitrate}$ independent of changes associated with denitrification (e.g. Studer et al., 2021). In addition, upwelled waters from dysoxic environments have lower N:P ratios due to N loss from denitrification. This may encourage biological $N_2$ fixation in the surface ocean, which brings in newly fixed N with lower $\delta^{15}N$ values (−1‰ to 1‰) (Sigman et al., 2009a, b; Ryabenko, 2013). These processes could modify the $\delta^{15}N$ of nitrate upwelled from dysoxic waters, especially in regions above and downstream of ODZ cores (e.g. Knapp et al., 2016; Wang et al., 2019b). Careful selection of paleoceanographic study sites and modelling studies constrained by analyses of modern $\delta^{15}N_{nitrate}$ from dysoxic environments is needed to separate influences from $NO_3^-$ consumption and $N_2$ fixation. Finally, various processes can influence the incorporation of the $\delta^{15}N$ signal into different organic N pools in the surface ocean and its preservation during sinking and burial. These processes are specific for each N archive and are discussed in the following sections.

## 7.2 Bulk sedimentary N isotopes

There is generally a good correlation between the upper-ocean $\delta^{15}N_{nitrate}$, sinking particulate $\delta^{15}N$, and surface sediment $\delta^{15}N$ ($\delta^{15}N_{bulk}$) in or near ODZs (Tesdal et al., 2013). As a result, $\delta^{15}N_{bulk}$ in marine sediments at or near an ODZ has been used to reconstruct changes in past water column denitrification and oxygen content, with most applications during the recent glacial/interglacial cycles (e.g. Altabet et al., 1995; Ganeshram et al., 1995; Galbraith et al., 2008, 2013). Interpretations of $\delta^{15}N_{bulk}$ records rely on the assumption that the total N preserved in the sediments represent the total N generated and exported out of the surface ocean. This assumption has been challenged in two main aspects (Robinson et al., 2012, and references therein). First, it has been widely observed that $\delta^{15}N_{bulk}$ is modified during sinking and incorporation into marine sediments. Sinking particles collected in the water column often show decreasing $\delta^{15}N$ with water depth (Altabet et al., 1991), which may be due to disaggregation of large particles into smaller particles, and thus infer a preferential loss of isotopically heavy N forms. Upon burial, microbial degradation processes may increase $\delta^{15}N_{bulk}$ in surface sediments with respect to sinking particles (by 2‰–5‰) (Altabet and Francois, 1994). Diagenetic effects on $\delta^{15}N_{bulk}$ have been re-

garded as small in ODZs because of low-oxygen content and high sedimentation rates, but this requires further validation. Second, $\delta^{15}N_{bulk}$ integrates the $\delta^{15}N$ signal of different N compounds in marine/terrestrial organic matter and clay-bound inorganic nitrogen. As such, it can be biased by the contribution of organic and inorganic N derived from terrestrial or distal marine (e.g. shelf) sources (Schubert and Calvert, 2001; Meckler et al., 2011). Thus, $\delta^{15}N_{bulk}$ is often used in combination with the total N content, carbon (C) to N ratio, and organic carbon isotope composition ($\delta^{13}C_{org}$) to quantify different endmember contributions using a simple mixing model. However, uncertainties using these calculations can be significant. Other efforts have been made to measure the $\delta^{15}N$ of total sedimentary N and clay-bound inorganic N separately to calculate the $\delta^{15}N$ value of the organic N ($\delta^{15}N_{org}$) (e.g. Schubert and Calvert, 2001). Despite these challenges in applying and interpreting $\delta^{15}N_{bulk}$, it remains one of the most used $\delta^{15}N$ archives in ODZs, especially in sediment cores with lower carbonate content. It can be analysed relatively quickly and easily and at a lower cost compared to other $\delta^{15}N$ archives.

## 7.3 Foraminiferal calcite-bound N isotopes

N isotopes in planktic foraminifera (FB-$\delta^{15}N$) have emerged as a novel approach to reconstructing past ocean oxygenation (Ren et al., 2009; Kast et al., 2019; Studer et al., 2021; Auderset et al., 2022; Wang et al., 2022b). When Rotaliid foraminifera build their chambers, they first produce an organic layer, upon which the calcite is precipitated (Branson et al., 2016). These organic sheets are mainly composed of proteins and polysaccharides, which are encased within the shells after calcification (Hemleben et al., 1989; Paoloni et al., 2023). Thus, foraminiferal-bound organic matter is protected by the calcite shells and less prone to diagenesis than bulk sediments. The amino-acid composition of the foraminiferal-bound organics of modern foraminifera appears to be very similar to that of fossil foraminifera from millions of years ago (King and Hare, 1972; Kast et al., 2019), suggesting there is little breakdown of more labile amino acids. Laboratory experiments have also shown that neither the N content nor its $\delta^{15}N$ vary with oxidative degradation, fossil dissolution, and thermal alteration, confirming the robustness of FB-$\delta^{15}N$ with respect to diagenesis (Martínez-García et al., 2022).

Planktic foraminifera incorporate N primarily by feeding on primary producers and zooplankton (Hemleben et al., 1989). Assimilation of $NO_3^-$ and $NH_4^+$ from the environment by photosynthesizing symbionts hosted in foraminifera could occur, but this is considered to be negligible in much of the oligotrophic oceans where the concentrations of both dissolved inorganic nitrogen (DIN) forms are low (Ren et al., 2012a, b). However, the symbionts could contribute to the internal N cycle by assimilating the recycled $NH_4^+$ excreted by the foraminiferal host. This process may reduce the over-
all ammonium leakage and lower the expected trophic en-
richment associated with it (Ren et al., 2012a, b). These pro-
cesses propagate the δ¹⁵Nnitrate signal through the food web
into the biomass of foraminifera. It has been shown that mod-
ern planktic FB-δ¹⁵N provide a robust record of δ¹⁵Nnitrate
in thermocline waters (e.g. Ren et al., 2009, 2012a), which
allows for the use of foraminifera from ODZ-influenced re-
gions for the reconstruction of past ocean suboxia (e.g. Kast
et al., 2019; Studer et al., 2021; Auderset et al., 2022; Wang
et al., 2022b).

7.4 Other potential archives for N isotopes in dysoxic environments

Diatom-bound δ¹⁵N (δ¹⁵N_db) is a potential yet unexplored proxy for past ocean oxygenation. Diatoms are siliceous phytoplankton that assimilate NO₃⁻ as their main N source with an isotopic offset of roughly 5‰ (Waser et al., 1998). During biomineralization, a fraction of that biomass is incorporated into their silica shell, termed frustule, which is thought to protect the organic matter from bacterial degradation and diagenetic alteration during sinking and burial in the water column and the sediment (Sigman et al., 1999). While the exact controls on the isotopic fractionation between diatom biomass N and frustule-bound N are not yet fully explored, δ¹⁵N_db has been shown to correlate with the δ¹⁵N of the NO₃⁻ source to the diatoms (Horn et al., 2011; Jones et al., 2022). Since diatoms thrive in areas of high nutrient supply, most paleoceanographic applications of the δ¹⁵N_db proxy have so far focused on the Southern Ocean and the sub-Arctic/Arctic oceans (Robinson and Sigman, 2008; Studer et al., 2012, 2015). While diatom opal also accumulates in the EEP upwelling region, diatoms are difficult to separate from radiolarians in those sediments, which can bias the δ¹⁵N_db signal (Studer et al., 2013; Robinson et al., 2015). Furthermore, incomplete NO₃⁻ consumption in EEP surface waters complicates the interpretation of δ¹⁵N_db as an oxygen proxy, as it is influenced by changes in both denitrification and NO₃⁻ assimilation. As such, δ¹⁵N_db has the potential to record past ocean oxygenation only in regions of complete surface NO₃⁻ consumption. Often, these oligotrophic environments are not hotspots for diatom accumulation on the seafloor. Nonetheless, a recent global compilation of opal flux records indicates that sedimentary opal concentrations reach > 10% in the northeastern Atlantic and equatorial Indian Ocean (Hayes et al., 2021). Those areas could be potential targets for future studies investigating past ocean oxygenation using diatom-bound N isotopes.

Most deep-water marine sediment records do not have
the temporal resolution necessary to capture the seasonal,
annual, or decadal climate variability required for di-
rect comparison with historical observations. In contrast,
reef-building, scleractinian corals are unique environmen-
tal archives that have been used extensively to reconstruct
climate variability during past centuries at high resolution

(Wang et al., 2016b). Typical growth rates in most coral
species used in paleoclimatic reconstructions are in the range
of ~1–2 cm per year, allowing for monthly sampling resolu-
tion in most cases. Methods have been developed to anal-
yse the δ¹⁵N of coral-skeleton-bound organic matter (CS-
δ¹⁵N), requiring as little as 5 mg of coral carbonate (Wang
et al., 2015). Modern ground truthing has demonstrated that
CS-δ¹⁵N follows NO₃⁻ assimilated across a wide range of
isotopic compositions and environmental settings (Wang et
al., 2014, 2016b). As a result, corals living close to the ma-
jor ODZs are ideal candidates for studying past changes in
water column denitrification and ocean oxygenation at high
temporal resolution. For example, a monthly-resolved CS-
δ¹⁵N record from the Oman margin shows values as high
as 10‰–11‰, apparently recording the ODZ-sourced NO₃⁻
signal in the Arabian Sea (Wang et al., 2016b). Despite their
slower growth rates, deep-sea corals (both scleractinian and
proteinaceous) can also be used to investigate the marine N
cycle and ocean deoxygenation in the past (McMahon et al.,
2015; Sherwood et al., 2014; Wang et al., 2014).

7.5 Case study: N isotopes in the eastern tropical Pacific

An increase in the overall denitrification rate may be ex-
pected to result from an increase in anoxic water volume as
a result of global warming and a decline in the overall ocean
oxygen concentration. Sedimentary δ¹⁵N records could help
constrain the various factors governing changes in water col-
umn denitrification rates and the oxygen content in the ocean.
Below we discuss a nitrogen isotope case study, focusing on
the dynamics of the eastern tropical Pacific ODZ across dif-
ferent time intervals and climatic backgrounds.

While previous Cenozoic records using δ¹⁵Nbulk show nearly no changes (Algeo et al., 2014), Auderset et al. (2022) and Kast et al. (2019) found a decline in FB-δ¹⁵N during prolonged warm periods, indicating a reduction in water col- umn denitrification and better ventilation of the Pacific ODZs (Fig. 7). Two possible mechanisms are proposed to explain the FB-δ¹⁵N trend: (1) a reduction in equatorial upwelling as a result of a reduced atmospheric pressure gradient during the warm periods and/or (2) a decline in the biological pump efficiency, resulting in less regenerated nutrients in the deep ocean. Both processes would lead to a better-ventilated EEP and reduction in the ODZs.

In addition, an FB-δ¹⁵N record from the eastern tropical
South Pacific (Wang et al., 2022b) has shown a significant
long-term increase in δ¹⁵N since the Miocene (Fig. 7). A
δ¹⁵Nbulk record from the California margin (Liu et al., 2008)
has also shown an increase since ~2.1 million years ago.
These records indicate expansion of the ODZs in the east-
ern tropical Pacific when the global climate had been cool-
ing since the mid-Miocene (Herbert et al., 2016; Westerhold
et al., 2020). Although decreasing sea surface temperatures
may have supplied more oxygen to the surface ocean since
the late Miocene, Wang et al. (2022b) argue that the rising oceanic nutrient content and resulting higher productivity appear to have overwhelmed the solubility effect and driven ocean deoxygenation over the past 8 million years.

On glacial/interglacial timescales, most published δ¹⁵Nbulk records from the eastern tropical Pacific generally show lower δ¹⁵Nbulk values during the last ice age compared to the Holocene warm period, which has been interpreted as lower water column denitrification and thus higher oxygen content in the water column (Ganeshram et al., 1995; Galbraith et al., 2004; Dubois et al., 2011). However, two recent FB-δ¹⁵N records from the EEP show similar δ¹⁵N values during the LGM and the Holocene, indicating comparable water column oxygen content during those periods (Studer et al., 2021) (Fig. 7). The difference between FB-δ¹⁵N and δ¹⁵Nbulk may be attributed to lower diagenetic alteration and/or higher foreign N input during ice ages, which may have altered the δ¹⁵Nbulk signal (Robinson et al., 2012; Studer et al., 2021). Using a box model, Studer et al. (2021) argue that multiple processes may have stabilized the oxygen content. A glacial shoaling of the Atlantic meridional overturning circulation, enhanced iron fertilization in the sub-Antarctic, and global cooling would have raised mid-depth oxygen content, whereas a more efficient biological pump would have led to an accumulation of regenerated nutrients and thus a decrease in deep Pacific oxygen content. This signal would have been mixed/upwelled into the mid-depth Pacific, leading to little net LGM-to-Holocene change in the eastern tropical Pacific ODZ extent (Hain et al., 2010; Studer et al., 2021). The FB-δ¹⁵N data are supported by the independent oxygen proxy I/Ca on planktic foraminifera (Hoogakker et al., 2018, Sect. 8) and challenge the previous views on reduced eastern tropical Pacific water column denitrification during ice ages based on δ¹⁵Nbulk (Ganeshram et al., 1995; Galbraith et al., 2004).

By combining different δ¹⁵N records from the eastern tropical Pacific across various timescales, a novel hypothesis has emerged that temperature-driven changes in mid-ocean oxygen content may not be the dominant control for ODZ evolution and changes in the water column denitrification rate. Instead, ocean circulation and biological activity are important additional controls (e.g. Robinson et al., 2014). These findings highlight the need for multiple N proxy applications in the same region or even the same sedimentary record, in combination with other oxygen proxies. In particular, we note an interesting recurring pattern in which δ¹⁵Nbulk records tend to be similar to the FB-δ¹⁵N when δ¹⁵N and the denitrification rate increase, but they often fail to record δ¹⁵N decreases across a climate event (Fig. 7 and references therein). As we generally assume better preservation conditions under low-oxygen conditions, the high δ¹⁵N values recorded by the bulk sediment during intervals of decreased FB-δ¹⁵N and increases in water column oxygen content could then be due to an increase in the diagenetic effect

on δ¹⁵Nbulk (Robinson et al., 2012). The δ¹⁵N difference between the bulk sediment and fossil-bound N may in turn provide interesting information on past changes in oxygen content in the water column and sediments. Finally, as physical and biological processes are both important for understanding the history of the ODZs, an oxygen-realistic biogeochemical model embedded with nitrogen isotopes would be crucial for advancing our interpretation of these δ¹⁵N records.

# 8 Foraminiferal trace elements

## 8.1 Introduction to proxy/geochemical system

In addition to sedimentary trace elements, foraminiferal trace elements can provide information about environmental redox conditions. In this section we focus on foraminiferal I/Ca, Mn/Ca, and U/Ca proxies, which are commonly used to track seawater redox conditions and relative dissolved oxygen concentrations. Specifically, these proxies are thought to track dissolved iodate (IO₃⁻), manganese (Mn(II)), and uranium (U(VI)), respectively. One frontier of the carbonate lattice-based proxies lies with determining the potential for their differential application to benthic and planktic foraminifera to quantify depth gradients in the water column. We discuss the proxy sensitivities within the context of the redox ladder to avoid confusion with the nomenclature (e.g. Canfield and Thamdrup, 2009). Below we discuss the distribution of IO₃⁻, Mn(II), and U(VI) in the water column and pore waters and their incorporation into foraminiferal calcite, and we evaluate their potential use as proxies for dissolved oxygen concentrations and seawater redox conditions (Fig. 8).

### 8.1.1 Seawater elemental cycling tracked via paleo-redox proxies

Multiple published seawater transects from the eastern tropical North Pacific and South Pacific document the geochemistry of Mn and I (Rue et al., 1997; Cutter et al., 2018; Rapp et al., 2020; Moriyasu et al., 2020, 2023). These studies demonstrate similarities and differences in seawater I and Mn speciation that can be exploited to define water column redox conditions. For example, redox thresholds for IO₃⁻ and Mn oxide reduction overlap (Fig. 8), and both display potentially rapid reduction kinetics, sluggish oxidation kinetics, and an important benthic component of accumulation within ODZs (Cutter et al., 2018; Rue et al., 1997). Specifically, in ODZs, dissolved iodate (IO₃⁻) is reduced to dissolved iodide (I⁻) at and below the oxycline (Fig. 8, areas labelled b). Estimates of the iodate oxygen reduction threshold range from up to 50 µmol kg⁻¹ (Lu et al., 2020) to potentially less than 1 µmol kg⁻¹ (Hardisty et al., 2021b). Thresholds less than 1 µmol kg⁻¹ may be possible, but iodine cycling has yet to be evaluated alongside oxygen sensors with sub-micromolar detection limits (e.g. Thamdrup et al., 2012). Regardless, the
<figure><img src="image_9.png" /><figcaption>**Figure 7.** Compilation of proxy reconstructions of equatorial Pacific anoxia from (FB)-δ¹⁵N. From left to right: the time spans 65–5 Myr, 5–0 Myr, 35–0 kyr, and 1860–2010 CE and the simulated global ODZ volume for RCP2.6 and RCP8.5 projections for 2020–2100 from IPCC AR5 (Bindoff et al., 2019). (a) Sea surface temperature compilation by Hansen et al. (2013). (b) Analyses of δ¹⁵N in bulk sediment (yellow) and foraminifera-bound (FB) organics (pink) in Pacific sedimentary archives. δ¹⁵N is used as a qualitative oxygen-sensitive proxy due to the strong isotopic fractions of denitrification by ODZ-dwelling bacteria. Data were compiled from the following sources: ODP 1209 (Kast et al., 2019); DSDP 598 (Wang et al., 2022b); ODP 872 (Auderset et al., 2022); ODP 1012 (Liu et al., 2008); ME24 and ME27 (Studer et al., 2021; Dubois et al., 2011); PES, SOL, and StaMON (Deutsch et al., 2014); the global Paleocene–Miocene compilation from Algeo et al. (2014); and the Paleocene–Eocene Thermal Maximum (PETM; Moretti et al., 2024).</figcaption></figure>

lowest IO₃⁻ has been demonstrated to nearly exclusively occur in low-oxygen waters, thus defining a threshold for “iodi-nous” conditions (Fig. 8) (Hardisty et al., 2021b).

At similar water column depths to where IO₃⁻ reduction occurs in ODZ cores (i.e. oxygen minima), dissolved Mn(II) also begins to accumulate (Fig. 8). Mn(II) is formed through the reduction in suspended manganese oxyhydroxide (III, IV) minerals driven by organic matter remineralization and/or oxidation of reduced components (e.g. dissolved iron, sulfides, and ammonium; Burdige, 1993; Thamdrup et al., 1994a, b; Kristiansen et al., 2002). Potential oxygen thresholds for Mn accumulation have been evaluated using switchable trace oxygen sensor (STOX) microsensors, which indicate that Mn(II) may be re-oxidized at oxygen levels as low as 100 nM (Clement et al., 2009). Ultimately, both Mn(II) and I⁻ are well-defined features in ODZs, and these are increasingly used alongside NO₂⁻ (product of nitrate reduction) maxima to define ODZ cores and so-called “functional anoxia”, which are zones where these and potentially more reducing metabolisms such as SO₄²⁻ reduction may occur

within microniches (e.g. Canfield et al., 2010; Raven et al., 2018).

Uranium cycling is a useful comparison to I and Mn cycling in that it is sensitive to more reducing ferruginous conditions (Fig. 1). U(VI) is dissolved in oxygenated seawater as uranyl carbonate (U(VI)) complexes and behaves conservatively (Ku et al., 1977; Chen et al., 1986; Dunk et al., 2002; Not et al., 2012). Uranium has a long residence time in seawater (~400 000 years), and therefore its concentration is homogenous and relatively constant on timescales of 100 000 years (Dunk et al., 2002). While U exhibits conservative behaviour in oxic settings, in oxygen-deficient/ferruginous seawater conditions, characteristic of some ODZ bottom waters, U(VI) is reduced to insoluble U(IV) and may precipitate on settling marine particles. Some of these U-containing particles ultimately reach the sediment to be either scavenged or re-oxidized to a soluble uranyl–carbonate complex at the sediment–water interface (Anderson et al., 1989). Under sufficiently reducing bottom and pore water redox conditions
<figure><img src="image_10.png" /><figcaption>**Figure 8.** Illustration of chemical profiles and key processes (a–i) affecting I/Ca, Mn/Ca, and U/Ca in foraminifera oxygenation proxies in a theoretical transect through an ODZ from (a) a deep-ocean site with overlying ODZ to (b) an intermediate-depth ODZ-adjacent site and (c) a shallow-water site with sediments in the ODZ. Based on data and concepts from Anderson et al. (1989), Thamdrup et al. (1994a, b), Morford et al. (2009), Singh et al. (2011), Koho et al. (2015), Moriyasu et al. (2020), and Singh et al. (2023). Note that when we refer to shallow water we mean the continental shelf, for intermediate waters we mean the continental slope, and for deep water we mean the abyssal plain.</figcaption></figure>

(dysoxic to anoxic), U will precipitate in the sediments as uraninite (U(IV)O₂(s); Fig. 8, areas labelled i).

Reduction in IO₃⁻, Mn oxides, and U(VI) takes place through bacterial catalysation. Reduction in IO₃⁻ has been associated with IdRA genes (Reyes-Umana et al., 2022). Soluble U(VI) is most likely reduced to insoluble uraninite U(IV) by iron-reducing bacteria (e.g. Lovley et al., 1991). Abiotic reduction is important as well. I and Mn are rapidly reduced via abiotic reactions with Fe and sulfide (Luther, 2023). Because ferruginous and sulfidic conditions are common in modern sediments, rapid reduction in IO₃⁻, Mn oxides, and U(VI) creates large gradients that can control diffusive fluxes to and from seawater. For IO₃⁻ and Mn oxides in sediments, rapid reduction creates elevated [I⁻] and [Mn(II)] – which

can be further exacerbated for [I⁻] due to organic matter remineralization (Kennedy and Elderfield, 1987a, b) – driving benthic fluxes into the overlying seawater (Fig. 8, areas labelled c and g). These fluxes are large within ODZ water columns, where low-oxygen concentrations enable the persistence of the reduced I⁻ (e.g. Cutter et al., 2018; Moriyasu et al., 2020) and Mn(II) (e.g. Froelich et al., 1979; Sundby and Silverberg, 1985; Metzger et al., 2007; Mouret et al., 2009) (Fig. 8). U behaves the opposite, where the formation of insoluble U(IV) under ferruginous conditions removes U from pore waters. This causes a concentration gradient to form between high-[U] overlying bottom waters and low-[U] pore waters and leads to a diffusive flux of seawater U into the sediments (Barnes and Cochran, 1990; Klinkhammer and
Palmer, 1991), causing sediment authigenic U (aU) enrich-
ment at a rate established by the diffusive flux (Fig. 8, area
labelled i).

8.1.2 Incorporation: how, when, and where are
elements incorporated?

The I/Ca, Mn/Ca, and U/Ca proxies track iodio-
nous, manganous, and ferruginous conditions, respectively
(Fig. 8). Below we consider what is known for each proxy.
Importantly, recent studies indicate that, whether applied to
planktic or benthic foraminifera, each proxy may, at least
in part, reflect the geochemistry of bottom waters, but more
work is needed.

It has been demonstrated that only the oxidized I species,
IO₃⁻, can be incorporated into both abiotic calcite (Lu et
al., 2010; Zhou et al., 2014; Podder et al., 2017; Kerisit et
al., 2018) and dolomite (Hashim et al., 2022) (Fig. 8, in-
set labelled a). Therefore, the I/Ca of planktic and ben-
thic foraminiferal calcite has traditionally been used to infer
[IO₃⁻] as a proxy for changes in subsurface and bottom wa-
ter oxygen concentrations, respectively (Glock et al., 2014,
2016; Hardisty et al., 2014, 2017; Hoogakker et al., 2018; Lu
et al., 2010, 2016, 2020; Zhou et al., 2014, 2016). However,
the incorporation of IO₃⁻ has not been directly tested within
foraminifera, which may include vital effects not currently
recognized. A recent study found that I/Ca values of plank-
tic foraminifera sampled from plankton tows showed little
to no relationship with the dissolved [IO₃⁻] of ambient sea-
water (Winkelbauer et al., 2023). In fact, this work showed
that planktic foraminiferal I/Ca was about 10 times lower
in plankton tows compared to that in sediment core tops,
as would have been expected from abiotic calcite precipita-
tion experiments (Winkelbauer et al., 2023). Winkelbauer et
al. (2023) suggest that planktic foraminifera may gain iodine
during gametogenesis or post-mortem, either when falling
through the water column or through burial. Lu et al. (2023)
confirm that proxy data from plankton tows and sediment
core-top samples may not agree because of the complexity of
foraminiferal calcification and post-depositional overprints
in marine surface sediments. Thus, core-top and downcore
planktic foraminiferal I/Ca may be representative of an in-
tegrated IO₃⁻ signal from across the water column and sedi-
ment, instead of the depth that they occupy during their life
cycle.

While both benthic and planktic foraminiferal I/Ca data
from core-top samples support a relationship with bottom
and subsurface water dissolved oxygen (Fig. 9a), the species-
specific and/or mineralogical controls for incorporation of
IO₃⁻ into biogenic carbonates are still not well understood.
Globally, the highest core-top I/Ca values are found to
be ~9 µmol mol⁻¹ for planktic foraminifera in the Walvis
Ridge region (Lu et al., 2020) and ~22 µmol mol⁻¹ for ben-
thic foraminifera in the Little Bahama Bank region (Lu et
al., 2022), both from well-oxygenated environments. Assum-

ing the IO₃⁻ concentrations in these oxic waters range be-
tween 0.5 and 0.65 µmol L⁻¹, the partition coefficient (Kd)
for IO₃⁻ incorporation (Kd = [I/Ca]foram / [IO₃⁻]sw, with a
unit of [µmol mol⁻¹] / [µmol L⁻¹]) can range from 14 to 44
or even higher if the seawater IO₃⁻ concentration is lower
than 0.5 µmol L⁻¹. These Kd estimates are much higher than
those reported in abiotic calcite synthesis experiments (~10)
(Lu et al., 2010), suggesting a potential biological control on
the I incorporation in the calcite. The strong association of
iodine with organic heterogeneities in the calcite of benthic
foraminifera might be another challenge to consider within
future studies on foraminiferal I/Ca (Glock et al., 2019a).

Interpretations of redox conditions based on the
foraminiferal Mn/Ca proxy may be derived from cal-
cite lattice-bound Mn (Mn/Caforam) and Mn bound in
post-depositional authigenic coatings of foraminifera tests
(e.g. Barras et al., 2018; Chen et al., 2017). Foraminifera
can incorporate soluble Mn(II) into their calcite tests (Fig. 8,
inset labelled e). Barras et al. (2018) observed a linear
correlation between Mn/Ca<sub>sw</sub> and Mn/Ca<sub>foram</sub> for two
different species (*Ammonia* T6 and *Bulimina marginata*)
of benthic foraminifera in controlled laboratory conditions.
That said, it seems that the partition coefficient increases
when concentrations of Mn<sub>sw</sub> are lower than ~10 µmol L<sup>-1</sup>.
The genus *Ammonia* is a shallow water taxon (inner neritic
– intertidal), and it is unclear if its biology is relevant to
deep-water taxa. Because of the link between Mn(II) and
ODZs, Mn/Ca in benthic foraminifera has been linked
to dissolved oxygen in the bottom and/or pore waters of
their microhabitat (e.g. Klinkhammer et al., 2009; Koho et
al., 2015). Concentrations of benthic foraminiferal (lattice-
bound) Mn/Ca from low-oxygen environments and culture
experiments can range from 0.1 to > 150 µmol mol<sup>-1</sup> (Lea,
1999; Glock et al., 2012a; Koho et al., 2017; Barras et al.,
2018; Brinkmann et al., 2021) depending on oxygen in
bottom/pore waters. Planktic foraminiferal (lattice-bound)
Mn/Ca ratios from plankton tows and sediment traps also
seem to be linked to seawater oxygen, with higher Mn/Ca
relating to lower oxygen (Steinhardt et al., 2014; Davis et
al., 2023a). The advantage of foraminiferal calcite-bound
Mn/Ca ratios compared to a bulk sediment proxy such
as Mn/Al is that once precipitated, the Mn concentration
remains fixed in the foraminiferal shell and should not be
subject to diagenetic reduction or oxidation (Koho et al.,
2015; McKay et al., 2015).

Nevertheless, in the case of fossil tests, post-mortem Mn-
rich contaminant secondary coatings (e.g. Mn oxides or Mn
carbonate; Fig. 8, inset labelled d) may obscure the Mn/Ca
signal of the pristine calcite signal (Barker et al., 2003; Ni
et al., 2020). Authigenic Mn mineral formation can occur
on the outside and/or inside of the foraminiferal tests and
pores. Recrystallization or banding within foraminiferal test
laminae can interfere with the application of this proxy for
reconstructing redox conditions when the foraminifera were
formed (e.g. Detlef et al., 2020; Ni et al., 2020). In addition
to the primary foraminiferal Mn proxy, several studies suggest authigenic foraminiferal U/Mn in coatings as a proxy for sedimentary post-deposit redox conditions (Gottschalk et al., 2016a; Chen et al., 2017; Detlef et al., 2020). The formation of Mn-rich authigenic carbonates potentially responds to the microbial activity in the pore water, which is linked to the sedimentary redox environment (Detlef et al., 2020; Ni et al., 2020).

The U/Ca proxy does not target carbonate lattice-bound U. Instead, this utilizes the formation of aU, which precipitates onto foraminifera tests buried in marine sediments, forming a U-rich (post-depositional) coating on their carbonate tests (Boiteau et al., 2012) (Fig. 8, area labelled i). The rate of authigenic enrichment is established by the U diffusive flux between overlying bottom waters and pore waters and follows similar dynamics to aU precipitation in bulk sediments (Boiteau et al., 2012). The diffusive flux, in turn, depends on how reducing the conditions are within the sediments (Barnes and Cochran, 1990; Klinkhammer and Palmer, 1991). Therefore, higher U/Ca concentrations are indicative of reducing oceanic bottom water conditions.

Concentrations of foraminiferal U/Ca can reach 300–700 nmol mol⁻¹ (Boiteau et al., 2012; Gottschalk et al., 2016a, 2020a; Skinner et al., 2019; Chen et al., 2017). This exceeds the foraminiferal lattice-bound [U], with shell matrix U/Ca ranging from ~1–23 nmol mol⁻¹ (Russell et al., 1994; Raitzsch et al., 2011; Boiteau et al., 2012; Yu et al., 2008; Chen et al., 2017). Therefore, lattice-bound U has a negligible impact on the measured U/Ca values of a diagenetically altered shell. Furthermore, the post-depositional accumulation of aU as foraminiferal authigenic coating means that any species can be measured, including planktic foraminifera, which tend to be much more abundant in open ocean settings.

### 8.1.3 Additional impacts on proxy values

Preservation and diagenetic effects on I/Ca ratios from biogenic calcite in sediments are currently unexplored, although it is well known that I/Ca in bulk carbonate is susceptible to diagenetic alterations, specifically declining values due to diagenesis in reducing IO₃⁻ free pore waters (Hardisty et al., 2017; Lau and Hardisty, 2022). Diagenetic impacts on local IO₃⁻ availability in pore fluids relative to overlying seawater may also make benthic foraminiferal signals particularly susceptible to recording diagenetic conditions. Specifically, a combination of excess I, related to I input from organic remineralization, and reducing conditions, which can impact IO₃⁻ availability, is possible. For example, I/Ca values as high as 20 µmol mol⁻¹ have been observed in infaunal foraminifera from the PETM. This may be due to higher IO₃⁻ near the sediment-water interface driven by oxic organic remineralization (Zhou et al., 2016; Kennedy and Elderfield, 1987a, b) or higher-than-modern total iodine concentrations in the seawater during the PETM. In the modern Peruvian

OMZ, shallow infaunal species (*Uvigerina striata* and *U. peregrina*) show I/Ca values ~1 µmol mol⁻¹ lower than epifaunal species (*Planulina limbata*) (Glock et al., 2014). If oxygen was the only control on pore water IO₃⁻, these lower I/Ca values in modern infaunal species would reflect lower IO₃⁻ concentrations in pore water, linking to rapid oxygen decrease within a few centimetres or millimetres of sediments. Thus, it remains unclear how pore water IO₃⁻ may influence the I/Ca of infaunal foraminifera. Attention should be given to possible contamination through organic-bound iodine in foraminiferal calcite (Glock et al., 2016, 2019a), which might significantly impact measured I/Ca ratios, thus requiring intense oxidative cleaning. Additional oxidative cleaning steps can result in considerably lower I/Ca ratios (Winkelbauer et al., 2021, 2023; see also Fig. 9a).

Like many other foraminiferal trace element proxies, several additional environmental parameters (e.g. temperature, salinity, and carbonate ion concentrations) may impact the elemental incorporation as well. In the Little Bahama Bank region where bottom water oxygen concentrations are similarly high between 150 and 200 µmol kg⁻¹, seawater temperature, salinity, and carbonate ion concentration show negative correlation with benthic foraminiferal I/Ca (Lu et al., 2022). It is thus speculated that benthic foraminifera may preferentially incorporate IO₃⁻ at lower temperature and/or lower salinity. Additionally, or alternatively, when ambient seawater has less carbonate ion availability, IO₃⁻ may be used by foraminifera as an alternative substitute for calcite structure. It is not clear whether the negative correlation between I/Ca and temperature could still be found under lower-oxygen bottom water conditions, as bottom water oxygen is often anti-correlated with temperature. Further studies are needed to disentangle such effects. Lastly, one epifaunal species, aragonitic *Hoeglundina elegans*, shows lower I/Ca values than observed in other benthic foraminifera (*Planulina limbata*, *Uvigerina peregrina*, and *Uvigerina striata*; Glock et al., 2014), suggesting different IO₃⁻ incorporation mechanisms for differing mineralogies (e.g. aragonite versus calcite) or an effect of different microhabitat preferences. Future work is needed to clarify these differences.

Factors such as carbonate chemistry, metabolic effects, ontogenetic effects, and species-specific effects could also have potential impacts on Mn incorporation into the foraminiferal test. Culture experiments with different partial pressures of carbon dioxide (pCO₂) levels show increased Mn/Ca of larger benthic foraminifera under high-pCO₂ conditions, which can be mainly ascribed to Mn speciation changes in seawater for Mn incorporation (Van Dijk et al., 2020). Mn incorporation can also be affected by Mg incorporation in hyaline species. Mg and Mn are coupled during foraminiferal calcification and are correlated on specimen and species level (Van Dijk et al., 2020). In the case of symbiont-bearing species grown under low-oxygen conditions, Mn mapping using electron probe microanalysis (EPMA) on the cross-section of the test highlighted that layers of calcite are en-
riched with Mn compared to ones grown under high-oxygen conditions (Van Dijk et al., 2019). This may be caused by the influence of the day/night (light/dark) cycles, meaning that symbiont activity (photosynthesis/respiration) or other diel shifts in physiology may directly or indirectly impact Mn concentration/speciation at the site of calcification. Dif-ferent chamber-to-chamber trends (such as between prolocu-lus and last chambers) for *Ammonia* and *Bulimina* species show ontogenetic effects (Barras et al., 2018, Brinkmann et al., 2023). Ontogeny-driven (life strategy) preferences may influence Mn/Ca in initial chambers (incl. proloculus) of *Bulimina*, as indicated by in-field observations (Brinkmann et al., 2023). Species-specific biomineralization processes and microhabitat effects could also impact Mn incorpora-tion in small benthic foraminifera (Koho et al., 2015; Bar-ras et al., 2018; Brinkmann et al., 2023; Groeneveld and Fil-ipsson, 2013; Groeneveld et al., 2018). Culture experiments demonstrate species-specific ontogenetic effects on Mn/Ca with opposite chamber-to-chamber trends in the last three chambers of *Ammonia* and *Bulimina* species (Barras et al., 2018). Diagenetic effects, including secondary mineral coat-ings, can significantly interfere with the Mn/Ca measure-ments of primary calcite under hypoxic/anoxic burial condi-tions (e.g. Detlef et al., 2020; Ni et al., 2020). The formation of inorganic carbonates with highly elevated Mn on the in-ternal and external surfaces of foraminiferal tests, and espe-cially in the pores, is difficult to eliminate through standard foraminiferal trace element cleaning procedures (Ni et al., 2020).

By contrast, post-depositional alterations form the basis of the U / Ca proxy. Authigenic U accumulated in sediments may be remobilized due to a deepening of the anoxic bound-ary, driven by an increase in bottom water oxygen, decreased C<sub>org</sub> flux, or bioturbation (McManus et al., 2005; Zheng et al., 2002b). Such a change in oxygen penetration depth could lead to a reversed gradient of pore (water) to bottom water [U], causing an efflux (remobilization) of dissolved U from the sediments back into the overlying bottom water, there-fore eliminating the primary U signal (McManus et al., 2005; Zheng et al., 2002a, b). It has been suggested that the U / Ca of foraminiferal coatings from marine sediment cores with high sedimentation rates ultimately records aU formed in steady state with bottom water and that the diagenetic aU loss is minimal (Gottschalk et al., 2016a; Jacobel et al., 2020).

8.2 How does the proxy relate back to oxygen?

Quantitatively relating proxies to oxygen content is an im-
portant goal for understanding paleo-redox evolution. This is
extremely challenging because each proxy has variable oxy-
gen thresholds and vital, mineralogical, and diagenetic ef-
fects that can distort its signal in the geologic record. Empir-
ical comparisons of plankton tow and core-top proxy values
to subsurface or bottom water oxygen can provide important
constraints on proxy-oxygen relationships applicable to the

geologic record. Towards this goal, Fig. 9 provides proxy-
oxygen syntheses, and below we provide both mechanistic
and empirical discussions to relate proxy records back to spe-
cific oxygen levels. For each of the proxies this means under-
standing at least two tiers of oxygen relationships: oxygen
values allowing for an initial change in the proxy from base-
line conditions and subsequent scaling of changing proxy
values to dynamic oxygen conditions.

I / Ca values recorded in planktic and benthic foraminifera
are lower in areas with lower subsurface and bottom wa-
ter oxygen, respectively (Fig. 9 and references therein). Of
the three foraminiferal proxies discussed here, I / Ca has the
best empirical constraints on oxygen thresholds, but at the
same time the mechanistic understanding of factors driving
these relationships remains unclear (see Sect. 8.1). Recent
studies provide two ways to interpret I / Ca relative to oxy-
gen content or redox conditions. The first is the simple pres-
ence/absence of carbonate-associated iodine and hence the
presence/absence of IO₃⁻. Iodide (I⁻) and IO₃⁻ have a simi-
lar redox potential, and IO₃⁻ is thus quantitatively reduced to
I⁻ in productive anoxic settings. This implies that the simple
presence of IO₃⁻ or carbonate-associated iodine may be in-
dicative of oxygen at some level (e.g. Hardisty et al., 2014).
We note that a global compilation of IO₃⁻ in ODZs demon-
strates that IO₃⁻ may persist when oxygen is below detec-
tion; however, this is interpreted to reflect sluggish reduction
in IO₃⁻, not in situ IO₃⁻ production (Hardisty et al., 2021a;
Moriyasu et al., 2020; Cutter et al., 2018). Additionally, as
discussed earlier, the CTD oxygen sensor detection limits,
which are typically near 1 µmol kg⁻¹, are currently a limita-
tion to understanding IO₃⁻-oxygen thresholds, as nmol kg⁻¹
oxygen levels may support active I redox cycles, as has been
demonstrated for N and Mn (Clement et al., 2009; Thamdrup
et al., 2012).

A major challenge in interpreting I / Ca data quantitatively
is that the oxidant(s) responsible for I⁻ oxidation to IO₃⁻ is
unknown. Indeed, unambiguous in situ IO₃⁻ production has
not been observed under normal marine conditions. Oxygen
is not directly responsible for IO₃⁻ formation, at least abioti-
cally, as demonstrated by the long residence time – estimated
to range from 40 to < 0.5 years – of I⁻ in oxygenated sea-
water (Chance et al., 2014). Some recent culture studies pro-
pose that superoxide or ammonia-oxidizing bacteria may be
responsible for catalysing IO₃⁻ production, but results have
yet to be confirmed in natural marine settings or in cultures
without iodine in excess of seawater (Hughes et al., 2021;
Li et al., 2014). While I⁻ oxidation is a limitation for mod-
els used to interpret mechanisms and distributions of ancient
IO₃⁻, constrained via I / Ca (e.g. Lu et al., 2018), IO₃⁻ re-
duction is clearly linked to declining oxygen, thus bolstering
proxy applications.

Beyond presence/absence, I / Ca values above and be-
low a threshold range may be used to define iodinous con-
ditions typical of ODZs. The iodinous framework allows
for I / Ca interpretations in the context of the redox lad-
<figure><img src="image_11.png" /><figcaption>**Figure 9.** (a) I/Ca in the core top (top 10 cm) and plankton tow planktic foraminifera versus minimum oxygen concentration in the water column. The data of Zhou and McManus (2023) were corrected for reductive cleaning. (b) I/Ca in epifaunal and infaunal benthic foraminifera versus bottom water oxygen concentration. (c), (d) Mn/Ca data from core-top benthic foraminifera compared to oxygenation in different settings. Note the differences in scale between (c) and (d). (e) OMZ and low-oxygen continental slope and (d) fjord, estuary, lake, and restricted basin settings. Note the two-tailed distribution from right to left, with (1) low Mn/Ca values where a deeper oxycline in pore waters results in less Mn²⁺ available at foraminiferal depth habitats, (2) high Mn/Ca values where low-oxygen bottom water prevents Mn²⁺ from leaving pore waters, and (3) low Mn/Ca where Mn²⁺ is lost to the water column with low-oxygen bottom water or precipitated as Mn carbonate. (f) U/Ca from planktic foraminifera deposited in sediments < 10 kyr BP compared to current bottom water oxygenation from the World Ocean Atlas 2018. Modified from Hu et al. (2023).</figcaption></figure>

der alongside other proxies, reflecting specific reduction potentials (Canfield and Thamdrup, 2009; Lau and Hardisty, 2022; Algeo and Li, 2020) (Fig. 1). I/Ca < 3 µmol mol⁻¹ can be related back to the IO₃⁻ range < 300 nM common to ODZ settings (Lu et al., 2016; Hardisty et al., 2021a; Lu et al., 2022). Lastly, specific oxygen thresholds have been recommended for the recognition of iodinous conditions. I/Ca values < 3 µmol mol⁻¹ have been demonstrated in benthic foraminifera with a bottom water oxygen concentration < 50 µmol kg⁻¹ (Fig. 8; Lu et al., 2022). I/Ca variations > 3 µmol mol⁻¹ are unlikely directly related to oxygen but instead likely reflect combinations of biologically mediated transformations during primary production and physical mixing and advection processes (Chance et al., 2014; Campos et al., 1996; Truesdale and Bailey, 2000; Hepach et al., 2020).

The application of Mn/Ca is mainly limited to environmental conditions characterized by trace amounts of bottom water oxygen. Whilst water column Mn cycling systematics are well understood, the direct relationship of Mn/Ca values to specific oxygen levels is restricted in comparison to I. Fundamentally, the highest Mn/Ca values are found in benthic foraminifera from manganous environments, allowing for high dissolved Mn²⁺ beneath intermediately oxic water columns that limit benthic Mn fluxes out of the sediments (Fig. 9c, d). This is because Mn/Ca tracks the reduced Mn endmember, contrarily to I/Ca, and because abiotic and benthic cycling can contribute to Mn(II) accumulation. For Mn, the diffusion of dissolved Mn from pore waters into bottom water, under prolonged anoxic conditions, prevents a linear relationship between Mn/Ca in foraminiferal calcite and bottom water oxygen concentrations (Koho et al., 2015). Recently, Brinkmann et al. (2023) found that an upper limit of around 130 µmol L⁻¹ exists in a fjord setting, above which the linear correlation between foraminiferal Mn/Ca and bottom water oxygen no longer exists, confirming earlier work by Guo et al. (2019) from the Yangtze River estuary. This proxy would also be difficult to apply in environments with very low dissolved Mn content, as is the case, for example, in the Santa Barbara Basin (Brinkmann et al., 2021). In the case of a depleted Mn pool in the sediment, the changes in Mn speciation according to oxygen concentrations would not be significant enough to be recorded in the foraminiferal shell. On the other hand, the vicinity of continental inputs or other sources of Mn into the sediment, independent of oxygen conditions, could also hamper the proxy robustness (Klinkhamer et al., 2009).

The Mn/Ca proxy is best related to relative oxygen levels, with the highest Mn/Ca encountered under hypoxic conditions. The Mn/Ca proxy will reflect the oxygen conditions in the microenvironment surrounding the foraminifera during calcification, i.e. bottom water, pore water, or water column conditions in the case of planktic foraminifera. Benthic foraminiferal species considered to be epifaunal or shallow infaunal should therefore better record bottom water condi-
tions than intermediate and deep infaunal taxa. Moreover, un-
der oxic conditions, epifaunal species would incorporate less
Mn. With decreasing bottom water oxygen concentrations,
the redox boundary would migrate towards the sediment-
water interface, and Mn incorporation would increase for
the shallow infaunal species. However, under anoxic/hypoxic
bottom water conditions, dissolved Mn diffuses out of the
sediment, resulting in less free Mn available for incorpora-
tion in the foraminiferal shell (e.g. Groeneveld et al., 2018).
Intermediate and deep infaunal taxa are already living at or
below the redox boundary and could migrate in the sediment
accordingly, potentially changing their microhabitat.

Similar to Mn/Ca, the U/Ca proxy is best related to relative increases/decreases in oxygen rather than absolute oxygen values. Higher U/Ca concentrations in foraminifera indicate more reducing sedimentary conditions, which are driven by low-oxygen concentrations in bottom and pore waters (Fig. 9e). However, these are driven by multiple processes of a physical or biological nature. Specifically, high organic C fluxes can lead to low-oxygen environments in pore waters and bottom waters, which ultimately cause the in situ precipitation of aU as foraminiferal coatings.

8.3 Description of analyses and resources required

Foraminiferal elemental analyses require careful species se-
lection, preparation, and cleaning and analytical procedures,
which can vary from element to element. For example, as
outlined in Table 3, the I / Ca and Mn / Ca paleo-redox prox-
ies specifically target lattice-bound I and Mn, while the
U / Ca paleo-redox proxy targets aU, meaning differential
cleaning procedures are required. In addition, while recent
advances have been made (Zhou et al., 2022; Cook et al.,
2022), the matrices used for analyses can vary between el-
ements, thus requiring careful consideration and, in some
cases, inhibiting multi-elemental analyses. Below, we pro-
vide an overview of cleaning and analytical procedures for
I / Ca, Mn / Ca, and U / Ca.

The foraminiferal cleaning procedures for both I / Ca and (calcite-bound) Mn / Ca are adapted from Mg / Ca protocols (Boyle and Keigwin, 1985; Barker et al., 2003). Samples are first gently crushed with cleaned glass slides to open all chambers, then cleaned by ultrasonication in deionized water to remove clays, next treated with a 10–20 min boiling-water bath in NaOH-buffered 1 % H₂O₂ solutions to remove organic matter, and lastly rinsed thoroughly with deionized water. It is recommended that samples with high organic matter should undergo additional oxidative cleaning steps (Glock et al., 2016; Winkelbauer et al., 2021). However, Mn and I analyses differ in that a reductive cleaning step is required for targeting lattice-bound Mn, as it is needed to remove authigenic Mn oxide coatings on the exterior of the test. For I / Ca, the reductive cleaning step is not required as the iodine content in Mn / Fe oxides is minimal (Zhou et al., 2014). Further, the reductive cleaning step has been demonstrated to cause a

systematic offset in I / Ca values and perhaps even Mn / Ca
(Fritz-Endres and Fehrenbacher, 2021), so it is not recom-
mended (Zhou et al., 2022). Notably, Mn carbonate coatings
are formed under reducing conditions and cannot be removed
with the existing cleaning techniques. Accordingly, increased
Mn / Ca values in foraminiferal calcite may either be part
of the test calcite itself or present as a coating. Diagenetic
contamination on the outer or inner surface of foraminiferal
tests like Mn carbonates and Mn-rich oxyhydroxides can be
identified with LA-ICP-MS, EMP, and XRF mapping. How-
ever, Mn overgrowth inside foraminiferal pores is difficult
to eliminate using LA-ICP-MS (Ni et al., 2020) due to the
material inside the pores being ablated through the whole
depth profile. We also need to consider that the Mn oxides
can be removed through a reductive cleaning step, whereas
authigenic Mn carbonates cannot be eliminated using stan-
dard solution-based cleaning methods, including the reduc-
tive cleaning step (Boyle and Keigwin, 1985).

In order to preserve authigenic foraminiferal coatings,
only a “weak chemical cleaning” protocol is used for U / Ca
foraminiferal analysis, whereby only the first step of clay
cleaning from the standard cleaning protocol from Barker et
al. (2003) is carried out. Cleaning experiments have shown
that the “Mg protocol” cleaning method (i.e. the addition
of oxidative cleaning) removes the authigenic coating and
therefore produces systematically lower U / Ca offsets in
comparison to clay cleaning only (Boiteau et al., 2012; Chen
et al., 2017).

While I / Ca, Mn / Ca, and U / Ca are all commonly mea-
sured via solution-based ICP-MS or ICP-OES, for the most
accurate I / Ca results, it is best to analyse samples in a
mildly basic solution (Cook et al., 2022). For example, while
carbonate dissolution via 1 %–3 % nitric acid (HNO₃) is suit-
able for other trace element / Ca ratios, samples for I / Ca
measurements are typically diluted in a 0.1 % or 0.5 % teri-
tiary amine, tetramethylammonium hydroxide (TMAH), or
ammonium hydroxide matrix to stabilize iodine in solution
before analyses (Winkelbauer et al., 2021). Other trace ele-
ments are commonly diluted in a 2 % HNO₃ matrix for anal-
ysis, which may be advantageous if the goal is to compare el-
emental data from the same foraminifera or to reduce the re-
quired sample mass and time/resources needed to separately
pick, clean, and measure two sets of foraminifera in order
to get separate I / Ca data. Importantly, a recent study cali-
brated a TMAH-nitric-based matrix that allows for simulta-
neous measurement of I / Ca, Mn / Ca, U / Ca, and a suite of
other trace elements (e.g. Li, B, Na, Mg, Al, Mn, Fe, Zn, Sr,
Cd, Ba, U, and Ca) (Cook et al., 2022). The drawback of si-
multaneously measuring a large suite of trace elements is that
the ideal cleaning and analysis procedures may vary from
element to element; some elements may thus be measured
under less-than-ideal circumstances. Differences in cleaning
or analytical procedures may also make comparison between
datasets measured using different protocols difficult.