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.