Cobalt is a nutrient to phytoplankton, but knowledge about its
biogeochemical cycling is limited, especially in the Pacific Ocean. Here, we
report sections of dissolved cobalt and labile dissolved cobalt from the US
GEOTRACES GP16 transect in the South Pacific. The cobalt distribution is
closely tied to the extent and intensity of the oxygen minimum zone in the
eastern South Pacific with highest concentrations measured at the oxycline
near the Peru margin. Below 200 m, remineralization and circulation produce
an inverse relationship between cobalt and dissolved oxygen that extends
throughout the basin. Within the oxygen minimum zone, elevated
concentrations of labile cobalt are generated by input from coastal sources
and reduced scavenging at low O
Cobalt is the least abundant inorganic nutrient in seawater, and its scarcity
may affect phytoplankton growth in certain regions
(Moore et al., 2013). In the high macronutrient waters
of the Costa Rica upwelling dome, for instance, Co and iron (Fe) amendments
to surface seawater increased phytoplankton production more than Fe alone,
promoting growth of the cyanobacterium
The GP16 transect in the tropical South Pacific. Red circles
indicate sampling stations. Dissolved oxygen at a depth of 300 m from the
World Ocean Circulation Experiment dataset is plotted in blue and 10
Biological cycling of dissolved cobalt (dCo) is apparent in vertical profiles, showing uptake and export in the surface and regeneration in the thermocline (Bown et al., 2011; Dulaquais et al., 2014a; Noble et al., 2012). While dCo in the euphotic zone can be entirely bound by strong organic ligands, a substantial portion of subsurface dCo is unbound and labile (10–50 %; Bown et al., 2012; Ellwood and van den Berg, 2001; Saito and Moffett, 2001; Saito et al., 2005) and therefore vulnerable to scavenging (Moffett and Ho, 1996). The similar ionic radii and redox potentials of cobalt and manganese (Mn) cause dCo to be actively incorporated into bacterially produced Mn oxides, which sink from the water column and accumulate in marine sediments (Cowen and Bruland, 1985; Moffett and Ho, 1996; Swanner et al., 2014). Below the euphotic zone, the persistence of labile dissolved cobalt (LCo) throughout the Atlantic indicates that scavenging of dCo, unlike Fe, is slow (Noble et al., 2012). On timescales of ocean circulation, however, scavenging is responsible for decreasing dCo concentrations with depth and for the low ratio between dCo and macronutrients in deep waters relative to phytoplankton biomass (Moore et al., 2013). As these deep waters are repackaged into thermocline water masses and eventually brought to the surface (Sarmiento et al., 2011), the upper ocean would become depleted in cobalt – as well as other hybrid metals like Fe and Mn – without external sources that keep pace with scavenging (Bruland and Lohan, 2003; Noble et al., 2008).
Yet the nature of marine cobalt sources is uncertain. In zonal sections of
the North and South Atlantic, sources appear to be concentrated along
continental margins (Noble
and Saito, 2016; Noble et al., 2012). In the western Atlantic, dCo
concentrations exceeding 100 pM were associated with the flow of Upper
Labrador Seawater, likely gained through intense sediment resuspension along
the shelf or input prior to subduction (Noble and Saito, 2016). dCo in
fresh and estuarine waters can be 100–1000 times greater than seawater
(Gaillardet et al., 2003; Knauer et al., 1982; Tovar-Sánchez et al., 2004), and Co is
less prone to flocculation in estuaries than other metals (Sholkovitz and
Copland, 1982). Terrigenous inputs from the American continent can be
clearly seen in lower-salinity surface waters influenced by the Gulf Stream
(Noble and Saito, 2016; Saito and Moffett, 2002) and Amazon discharge
(Dulaquais et al., 2014b).
Yet in both the North and South Atlantic, a much larger dCo plume was
associated with the oxygen minimum zones (OMZs) along the Mauritanian and Namibian
coasts (Noble and Saito, 2016, Noble et al., 2012). Although these
waters are not anoxic, the dCo plumes imply that O
To date, sectional datasets for dCo have been confined to the Atlantic, and,
as such, our understanding of cobalt cycling may be biased by the dominant
processes occurring there. In comparison, the South Pacific receives
considerably less dust deposition and river input
(Mahowald et al., 2005; Milliman and
Farnsworth, 2011) but hosts a much larger and more reducing oxygen minimum
zone. Surface transects off Peru and the Costa Rica Dome suggest a large
source from upwelling
(Ahlgren et al., 2014;
Saito et al., 2004, 2005); however, profiles in the tropical Pacific are
sparse
(Noble et al., 2008; Saito et al., 2014). We measured the concentration of
dissolved cobalt and labile dissolved cobalt in over 750 samples collected
onboard the 2013 US GEOTRACES GP16 expedition across the South Pacific along
12
Sampling for GP16 was conducted with a 24-position trace metal clean titanium
rosette attached to a nonmetallic Kevlar cable designed for the US
GEOTRACES program (Cutter and Bruland, 2012).
An additional sample was collected from a surface towfish at each station.
Subsamples were collected in a Class-100 sampling van from 12 L Go-Flo
bottles (General Oceanics) and passed through 0.2
dCo and LCo were measured using a cathodic stripping voltammetry (CSV)
method optimized for organic speciation by Saito and Moffett (2001). This method relies on the complexation of inorganic Co species by a
strong synthetic ligand, dimethylglyoxime (DMG, K
Triplicate scans of the seawater sample were followed by four standard
cobalt additions (25 pM per addition), and the slope of their linear
regression (mean
dCo analyses were conducted after a 1 h UV oxidation procedure to remove
strong organic ligands that prevent DMG from binding Co. UV digestion was
performed in 15 mL quartz glass tubes using a Metrohm 705 UV digester
(Metrohm USA). Temperature was maintained below 20
LCo was measured after more than 8 h incubation of 11 mL of seawater
with 400
All bottles and sample tubes were soaked for more than 1 week in the
acidic detergent Citranox, rinsed thoroughly with 18.2 M
To determine reagent blanks, Co-free seawater was generated by treating
UV seawater with cleaned Chelex-100 beads. The seawater was then UV digested
a second time to remove any ligands leached during Chelex treatment. Any dCo
measured in the Chelexed seawater derives from addition of Co from
analytical reagents. The mean blank for at-sea analysis was consistently
low: 3.7
Signal processing of voltammetry scans. Varying instrumental noise
imprinted negative current excursions during measurement and necessitated
data smoothing to correctly measure the Co(DMG)
To accommodate a greater number of samples, our previous workflow
(Noble et al., 2008) was modified to incorporate fully automated sample analyses
using the Metrohm 858 Sample Processor autosampler. All measurements were
performed using an Eco-Chemie
Tubes containing 11 mL seawater, DMG, and EPPS were inverted several times
and placed onto a sampling rack where 8.5 mL of the mixture was dosed into
the Teflon analysis cup. Then, 1.5 mL of 1.5 M sodium nitrite was added directly
to the sample cup. Samples were purged with high-purity N
We noticed a decrease in sensitivity of preserved samples relative to those
analyzed at sea, possibly caused by an increase in the sample pH during
storage. Sensitivity was restored by doubling the concentration of our
buffering agent, EPPS, in the sample to a final concentration of 7.6 mM. We
tested a broad range of EPPS additions in UV seawater and found the cobalt
concentration unchanged, while the variance between triplicate scans was
reduced markedly by the increase in sensitivity (data not shown). We
tentatively attribute this decrease in sensitivity in preserved samples to
CO
Analyses conducted at sea were characterized with a mild to moderate
electrical interference that mandated additional processing before peak
height could be reliably measured (Fig. 2). We opted for a simplified least
squared fitting routine included in the NOVA software package that conducts
a 15-point weighted moving average – equivalent to a 36.9 mV window –
according to a second-order polynomial. This method did not distort
measured peak height when noise was low (Fig. 2a, b). A small fraction of
scans (
Subsequent analyses in the laboratory at Woods Hole were able to remove this
signal by increasing the current sampling step from 2.46 mV (341 points
between
All data reported in this manuscript have been submitted to the Biological
and Chemical Oceanography Data Management Center (BCO-DMO,
Profiles of dissolved cobalt (dCo, closed circles), labile
dissolved cobalt (LCo, open circles), and O
Dissolved oxygen
Coupling between dissolved cobalt with O
Dissolved cobalt
Blanks and standards used during analyses.
Because acidified community reference materials such as the SAFe standards
require a delicate neutralization to pH 7.5–8 prior to analysis, a large
batch of UV oligotrophic seawater was generated prior to the cruise and used
to assess instrument performance during at-sea analysis. This consistency
seawater standard was run roughly three times per week, as were blanks, and
values were stable over several reagent batches for the duration of the
cruise (4.5
Particulate material collected from Go-Flo bottles was filtered onto
acid-cleaned 0.45
We report 680 determinations of dissolved cobalt (dCo) and 783 determinations of labile dissolved cobalt (LCo) measured at sea, onboard the GP16 expedition in October–December 2013, as well as an additional 140 measurements of dCo measured from preserved samples on land. In this section, we describe the distributions of dCo, particulate cobalt (pCo), and LCo in the South Pacific Ocean.
Throughout the GP16 transect, nutrient uptake and scavenging result in a hybrid-type profile for dCo (Fig. 3), similar to dCo profiles from the Atlantic (Bown et al., 2011; Dulaquais et al., 2014b; Noble et al., 2012; Noble and Saito, 2016) and North Pacific (Ahlgren et al., 2014; Knauer et al., 1982; Saito et al., 2014). dCo ranged from < 3 pM (below detection) in the South Pacific gyre (e.g., stations 23, 36) to 210 pM beneath the oxycline near the Peru margin (Station 1). In the deep Pacific, concentrations fell between 20 and 40 pM but increased slightly at deepest stations below 4500 m. These values are much less than those observed in zonal transects surveying the North and South Atlantic (Noble et al., 2012; Noble and Saito, 2016) but are similar to measurements in the Southern Ocean (Bown et al., 2011), indicating that dCo is scavenged in the deep ocean along meridional overturning circulation. Below 3000 m, dCo is somewhat lower east of the East Pacific Rise (EPR) and matches less oxygenated, older waters than in the western portion of the transect (Fig. 4). While many profiles west of the EPR show considerable variation between 2000 and 3000 m, suggestive of hydrothermal influence, the range is small (< 10 pM) relative to background concentrations (30–40 pM) and unlike the 50-fold excess of hydrothermal dFe and dMn above background seawater measured at Station 18 (Resing et al., 2015).
dCo peaks in the mesopelagic zone, typically between 300 and 500 m. Towards the Peru
shelf, this maximum shoals and increases, following the position and
intensity of the OMZ (defined here as < 20
All profiles show a surface or near-surface minimum that indicates
biological uptake and export. As a result, dCo is well traced by dissolved
phosphate, PO
The surface minimum in dCo is mirrored by a near-surface maximum in
particulate cobalt (pCo) from biological uptake throughout the GP16 section.
The distribution of pCo (Fig. 6c) resembles particulate phosphorus (pP),
chlorophyll, and other indicators of phytoplankton biomass. Very high pCo
(> 10 pM) was measured in the highly productive waters in the
Peru upwelling ecosystem, while lower concentrations (2–4 pM) were found in
oligotrophic surface waters. West of 100
dCo can be bound by extremely strong organic ligands that affect its
reactivity (Ellwood and
van den Berg, 2001; Saito and Moffett, 2001). These ligands may be composed
of degradation products of the cobalt-bearing cofactor vitamin B
In the ETSP, the distribution of LCo is similar to that of dCo (Figs. 3, 4, 6).
Except for samples from the upper 50 m, dCo and LCo form a linear
relationship (
In the deep Pacific (more than 3000 m), where dCo is low, LCo is
undetectable. LCo remains low (< 15 pM) in the mesopelagic zone, except
where the OMZ is most intense (Fig. 4). Within the OMZ, LCo maxima coincide
with dCo maxima (stations 1–15), but further to the west these LCo maxima
are much less pronounced and occur deeper than dCo maxima (Fig. 3). The LCo
plume from the OMZ also extends deeper (below 2000 m) than the corresponding
dCo plume (< 2000 m), suggesting that remineralization and scavenging
affect these quantities in different ways. Slight secondary maxima between
1500 and 2000 m (10–15 pM) appear in the center of the section on
The most striking aspect of the dCo distribution in the ETSP is the very
high concentrations present in the OMZ (Figs. 3–5). Similar distributions
have been observed in both the North and South Atlantic, where > 100 pM dCo plumes corresponded to low-oxygen waters underneath the Benguela
and Mauritanian upwelling systems
(Noble et al., 2012; Noble and
Saito, 2016). In the North Pacific, profiles from the Costa Rica Dome
(Ahlgren et al., 2014), the California margin
(Biller and Bruland, 2013; Knauer et al., 1982), and the central Pacific along
155
In the following, we describe the oceanographic processes that lead to
elevated dCo and LCo concentrations in low-oxygen waters (Sect. 4.1). On the basin scale, the combined effects of remineralization and circulation
link dCo with O
In the ETSP, tight coupling between dCo and O
In the deep ocean, near-conservative mixing of low-O
In the upper 200 m, dCo is not well coupled with O
In the upper ocean (0–200 m), dCo is linearly related to PO
Transition in dCo cycling at the OMZ boundary in the upper South
Pacific thermocline.
A separate nutrient-like dCo
In the eastern margin, the surface and mesopelagic dCo
The isopycnal dCo
Crossing the anoxic–oxic transition at 100
Redox control of Co and Mn scavenging. Within mesopelagic waters
(
Profiles of dissolved cobalt (dCo, black), PO
The stimulation of cobalt scavenging across the anoxic–oxic transition at
100
The strong covariation between high dCo and low O
Cobalt and Mn in the Peru shelf OMZ (GP16 stations 1–5, O
Positive correlations between dCo, LCo, and dMn within the OMZ on the Peru
shelf reflect a shared source (Fig. 11). The slope of the LCo
Water column observations of a large dCo source are also mirrored in the
depleted Co contents of continental shelf sediments along the Peru margin. A
survey of continental shelf sediments underlying the Peru OMZ found low
Co
Co
The Co
In contrast to depleted Co along the South American shelf, the Co
It is likely that oxidizing conditions in the water column and surface
sediments limit the release of cobalt on the western margin, leading to
crustal Co
Altogether, the accumulating evidence that the oceans' major OMZs harbor dCo plumes (Noble et al., 2012; Alhgren et al., 2014; Noble et al., 2016) indicates a strong chemical connection between the efficiency of dCo sources and local redox conditions where these OMZs interact with the continental margin. While reductive dissolution is the most likely mechanism at play, whether this process predominantly occurs in estuaries, the OMZ water column, or margin sediments is presently unclear but will ultimately dictate how the dCo source (and the resulting OMZ plumes) are affected by climate-driven changes in the size and intensity of OMZs (e.g., Scholz et al., 2011, 2014). As such, there is a significant need for future experimental and field studies that address the redox sensitivity of dCo and other metal fluxes from coastal environments.
Can a terrigenous cobalt source account for the observed OMZ plume? Because
lithogenic sediments along the Peru margin are delivered primarily by rivers
(Scheidegger and Krissek, 1982),
we can estimate a dCo flux to OMZ waters as the product of the fluvial
sediment delivery to the continental shelf and the difference in Co
The extent to which the coastal flux and dCo inventory are in agreement
depends on the residence time of OMZ waters. Models and CFC distributions
from the World Ocean Circulation Experiment (WOCE) dataset imply an approximately decadal recirculation time in OMZ waters
relative to mesopelagic gyre circulation in the ETSP
(Deutsch et al., 2001, 2011). Integrating our terrigenous
Co flux estimate over 10 years yields an expected concentration of 120–230 pM within the OMZ. This is of similar magnitude but greater than the
concentrations measured in the GP16 dataset (mean of 100
We can compare the calculated sedimentary flux to an expected flux from
aerosol dust dissolution. Aeolian deposition is extremely low over the South
Pacific basin (Mahowald et al., 2005),
except immediately offshore of Peru, where dust from the Altiplano interacts
with the prevailing northward winds
(Prospero and Bonatti, 1969; Scheidegger and Krissek, 1982). Model results
(Mahowald et al., 2005) suggest that
deposition does not exceed 0.5 g m
Hydrothermal venting along the EPR provides a major
source of dFe and dMn to the deep South Pacific (Resing
et al., 2015) where nanomolar concentrations of both metals were measured
between 2000 and 3000 m at the ridge crest and concentrations exceeded
background values for several thousand kilometers westward. In contrast, dCo
concentrations are only slightly elevated at the ridge crest (Station 18,
Fig. 3), reaching 36 pM at 2400m (against a background of
Profiles from Station 18 at the East Pacific Rise ridge crest at
113
However, both dCo and LCo maxima are offset from dFe and dMn plumes. At
Station 18, dFe and dMn peak at 2500–2600 m. At this depth, LCo is
undetectable and dCo values are at – or slightly lower than – background
levels (Fig. 12b), suggesting that Mn and/or Fe scavenging in the heart of
the hydrothermal plume has removed most of the hydrothermal Co from the
water column before being transported away from the ridge crest. Indeed, Co
is strongly associated with Mn phases in near-axis metalliferous sediment in
the EPR at 14
The combination of eastern boundary upwelling and a continental source
produces large dCo gradients across the surface of the South Pacific Ocean
(Fig. 6a). Westward, decreasing surface dCo results from phytoplankton
uptake and export, reflected in strong correlations with PO
While productivity in the South Pacific is thought to be limited by scarcity
of iron and nitrogen (Moore et al., 2013;
Saito et al., 2014), the extremely low dCo measured here implies that it may
be important as well. Because marine cyanobacteria such as
Schematic cross section of the cobalt cycle in the eastern tropical South Pacific. Black arrows describe idealized physical
circulation, showing upwelling near the Peru margin, advection westward, and
subduction in the South Pacific gyre. Biological Co export is shown in the
red-striped arrows, and solid and dashed red arrows show remineralization and
scavenging, respectively. The margin source is shown as a red-outlined
arrow. These vectors are also plotted on idealized oxygen and phosphate
axes, using the same color scheme, to show how these processes appear in
Co
Unlike its near uniform relationship with dCo in the underlying OMZ, LCo
measured during GP16 is low relative to dCo in the surface ocean (0–50 m),
especially along the Peru margin (Figs. 7a, 13c). This might result either
from microbial production of cobalt ligands – as observed in a
Spatial and temporal variability of margin dCo sources may ultimately affect
carbon flow through the Peru upwelling ecosystem. Considering the very low
dissolved Zn in surface waters during GP16(< 100 pM east of
90
A schematic of the cobalt cycle in the eastern tropical South Pacific Ocean
– and how these processes lead to covariation of dissolved cobalt with
O
The basin-scale association between high dCo and low O
Ultimately, the dCo inventory in the South Pacific – and its availability
to phytoplankton – may be changing considerably as the size of the
OMZ fluctuates. Recent warming and stratification appear to have expanded
the volume of low-oxygen waters in the tropics
(Stramma et al., 2008, 2010). As such,
dCo inventories may increase as lower O
All data can be accessed on BCO-DMO (Saito, 2016; Sherrell and Twining, 2016). dCo and LCo databases can be downloaded
at
N. J. Hawco, D. C. Ohnemus, and J. A. Resing participated on the EPZT cruise. N. J. Hawco measured dCo and LCo. J. A. Resing measured dMn; D. C. Ohnemus and B. S. Twining measured particulate Co, Mn, and P. N. J. Hawco and M. A. Saito prepared the manuscript with contributions from all authors.
We thank the captain and crew of the RV