1. Introduction
Oceanic distributions of dissolved Cr show nutrient-type behavior, generally resembling the primary macro- and micronutrients, though with muted surface depletions and enrichments at depth (e.g. Campbell & Yeats, 1981; Jeandel et al., 1987). Thermodynamic calculations predict that oxidized Cr(VI), a soluble oxyanion with low particle reactivity (Semeniuk et al., 2016), should account for all dissolved Cr in the oxygenated ocean (Elderfield et al., 1970). However, the reduced form, Cr(III), is regularly found at low levels in oxygenated seawater (≤ ~15% of total dissolved Cr, e.g. Cranston & Murray, 1978; Cranston, 1983; Jeandel & Minster, 1987; Achterberg & van den Berg, 1997; Connelly et al., 2006; Janssen et al., 2020), and may constitute a major fraction of dissolved Cr in oxygen minimum zones (OMZs; e.g. Murray et al., 1983; Rue et al., 1997). Chromium(III) is relatively insoluble and readily adsorbs to mineral and organic surfaces (e.g. Cranston & Murray, 1978; Mayer et al., 1984). Therefore redox transformations likely play a central role in Cr biogeochemical cycling. Iron(II) is an important Cr(VI) reductant in natural aquatic systems (Pettine et al., 1998), and Mn oxides are a primary Cr(III) oxidant (van der Weijden & Reith, 1982; Milletto et al., 2021), especially in Mn-rich environments such as sediments (Oze et al., 2007).
The primary dissolved Cr removal processes in the ocean, both involving Cr reduction followed by particle scavenging, are removal in OMZs and biologically-mediated export (e.g. Scheiderich et al., 2015). Chromium removal in OMZs has been reported in the eastern tropical North Pacific (ETNP) (Murray et al., 1983; Rue et al., 1997; Moos et al., 2020) and the eastern tropical South Pacific (ETSP) (Nasemann et al., 2020), but this process may be more efficacious in anoxic shelf environments (cf. Moos et al., 2020; Nasemann et al., 2020). Correlations between Cr and particulate organic carbon in sinking particles (Connelly et al., 2006), Cr(III) adsorption onto marine phytoplankton (Semeniuk et al., 2016), and agreements between dissolved Cr deficits and productivity-based inferred removal (Janssen et al., 2020) support a Cr sink associated with biological export. However, the subtle gradients in [Cr] depth profiles suggest that release from biogenic particles is relatively minor (e.g. Scheiderich et al., 2015; Goring-Harford et al., 2018; Moos & Boyle, 2019; Rickli et al., 2019) meaning the contribution of biogenic Cr fluxes to concentrations in the ocean interior may be limited, and circulation may play an important role in shaping distributions (Connelly et al., 2006; Scheiderich et al., 2015; Goring-Harford et al., 2018; Rickli et al., 2019). Elevated bottom water [Cr] suggests surface to deep [Cr] gradients may be influenced by a flux from marine sediments (Murray et al., 1983; Jeandel and Minster, 1987; Achterberg & van den Berg, 1997) consistent with shallow pore water Cr enrichments on the California shelf due to release from organic matter (Shaw et al., 1990).
Chromium redox transformations are accompanied by isotopic fractionation. While fractionation patterns during Cr oxidation are variable (see discussion in Zink et al., 2010; Milletto et al., 2021), reduction consistently enriches light isotopes in Cr(III) (e.g. Wanner & Sonnenthal, 2013). The redox control on [Cr] and δ53Cr makes δ53Cr a potentially powerful tracer of paleoredox conditions, especially in early earth studies (e.g. Frei et al., 2009). However, application and interpretation of δ53Cr records require an accurate and mechanistic understanding of the process(es) that control Cr budgets and isotopic fractionation in terrestrial environments, the global ocean, and during incorporation of Cr into sediment records. The systematic relationship between seawater [Cr] and δ53Cr suggests that redox transformations likely control the global distribution of oceanic δ53Cr (Scheiderich et al., 2015; Goring-Harford et al., 2018; Moos et al., 2019; Rickli et al., 2019; Janssen et al., 2020; Moos et al., 2020; Nasemann et al., 2020) implying an effective isotope enrichment factor (ε) of approximately -0.7 ‰ associated with Cr reduction and removal. Recent δ53Cr and [Cr] data demonstrate fractionation associated with biologically driven Cr removal in the open ocean (Janssen et al., 2020). Similarly, Cr removal in shelf settings (Goring-Harford et al., 2018), as well as in OMZs (Moos et al., 2020; Nasemann et al., 2020) can result in isotopic fractionation. However, isotope fractionation factors for these removal processes remain poorly constrained.
To address uncertainties in the modern ocean Cr budget and the role of biogenic processes in seawater δ53Cr distributions, we present dissolved [Cr], δ53Cr and [Cr(III)] from shipboard incubations, new intermediate and deep water [Cr] and δ53Cr data from eight research expeditions in the Southern, Pacific and Atlantic Oceans, and pore water [Cr] data from the Tasman Sea.