235U in New Zealand’s fjords: Implications for reconstructing global paleoredox conditions in organic-rich sediments

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ScienceDirect Geochimica et Cosmochimica Acta 179 (2016) 89–109 www.elsevier.com/locate/gca

Trace metal cycling and 238U/235U in New Zealand’s fjords: Implications for reconstructing global paleoredox conditions in organic-rich sediments Jessica L. Hinojosa a,b,c,⇑, Claudine H. Stirling b, Malcolm R. Reid b, Christopher M. Moy a, Gary S. Wilson c b

a Department of Geology, University of Otago, PO Box 56, Dunedin 9054, New Zealand Centre for Trace Element Analysis and Department of Chemistry, University of Otago, PO Box 56, Dunedin 9054, New Zealand c Department of Marine Science, University of Otago, PO Box 56, Dunedin 9054, New Zealand

Received 8 June 2015; accepted in revised form 5 February 2016; Available online 9 February 2016

Abstract Reconstructing the history of ocean oxygenation provides insight into links between ocean anoxia, biogeochemical cycles, and climate. Certain redox-sensitive elements respond to changes in marine oxygen content through phase shifts and concomitant isotopic fractionation, providing new diagnostic proxies of past ocean hypoxia. Here we explore the behavior and inter-dependence of a suite of commonly utilized redox-sensitive trace metals (U, Mo, Fe, and Mn) and the emerging ‘‘stable” isotope system of U (238U/235U, or d238U) in New Zealand fjords. These semi-restricted basins have chemical conditions spanning the complete redox spectrum from fully oxygenated to suboxic to intermittently anoxic/euxinic. In the anoxic water column, U and Mo concentrations decrease, while Fe and Mn concentrations increase. Similarly, signals of past euxinic conditions can be found by U, Mo, Fe, and Mn enrichment in the underlying sediments. The expected U isotopic shift toward a lower d238U in the anoxic water column due to U(VI)–U(IV) reduction is not observed; instead, water column d238U profiles are consistent in fjords of all oxygen content, falling within previously reported ranges for open ocean seawater (d238U =
0.42 ± 0.07‰). Additionally, surface sediment d238U results show evidence for competing U isotope fractionation processes. One site indicates increased export of 238U from seawater to the underlying sediments (fractionation between aqueous seawater U and particulate sediment U, or DU(aq)
U(solid) =
0.25‰), consistent with redox-driven fractionation. Another site suggests potential U(VI) adsorption-driven fractionation, reflecting increased export of 235U from seawater to sediments (DU(aq)
U(solid) = 0.25‰). We discuss several potential factors that could alter d238U in waters and sediments beyond redoxdriven shifts, including adsorption to organic matter in waters of high primary productivity, reaction rates for competing processes of U adsorption and release, and isotopic constraints of U coming into the system from terrestrial environments. These potential complications should be understood and constrained through observations, experiments, and models before future application of d238U as a global paleoredox tracer can achieve its full potential. Ó 2016 Elsevier Ltd. All rights reserved.

1. INTRODUCTION ⇑ Corresponding author at: California Institute of Technology, Geological and Planetary Sciences, 1200 E. California Blvd., MC 100-23, Pasadena, CA 91125, United States. Tel.: +1 626 395 6111. E-mail address: [email protected] (J.L. Hinojosa).

http://dx.doi.org/10.1016/j.gca.2016.02.006 0016-7037/Ó 2016 Elsevier Ltd. All rights reserved.

Recent advances in mass spectrometry have allowed the stable isotope systematics of an increasing suite of elements, including many transition metals, to be quantified. This isotopic information, when combined with elemental

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concentrations, provides additional discriminatory power and the ability to track past and present redox conditions in the environment (e.g., Anbar and Rouxel, 2007; Lyons et al., 2009). In particular, the low-temperature, redoxdriven isotopic fractionation between 238U and 235U is emerging as a potentially powerful paleoredox tracer. Recent studies have shown U isotopic fractionation of up to 5‰ during reduction and oxidation in natural environments (e.g., Stirling et al., 2007; Weyer et al., 2008; Bopp et al., 2009; Romaniello et al., 2013; Murphy et al., 2014; Andersen et al., 2014; Holmden et al., 2015; Noordmann et al., 2015b), which is believed to occur due to varying nuclear volumes between isotopes of different mass, known as the ‘‘nuclear field shift effect” (Schauble, 2007). By this mechanism, the heavier species (238U) is preferentially concentrated in the lower U(IV) oxidation state that has a decreased electron density in the nucleus (Schauble, 2007). Therefore, when U is reduced in low-oxygen environments, a ‘‘heavy” 238U/235U isotopic signature is imparted to the reduced U(IV) species exported to particulate form. Conversely, the remaining aqueous U(VI) becomes isotopically ‘‘light.” However, our understanding of redox-linked U isotope systematics is still nascent. Recent studies investigating modern calibration of the 238U/235U redox proxy have shown contrasting results. In the modern permanently anoxic Black Sea (Eckert et al., 2013) and Kyllaren fjord basins, 238U is preferentially exported from the waters to the underlying sediment, leaving an isotopically ‘‘light” water column enriched in 235U relative to open ocean seawater (Andersen et al., 2014; Noordmann et al., 2015b), which is the manner expected for redox-driven shifts under equilibrium conditions based on ab initio model calculations (Abe et al., 2008). In such settings, removal of up to 40% of aqueous U due to reduction has been observed (Andersen et al., 2014), which will lead to shifting U isotopic values in the water column and underlying sediments. A suite of studies from experiments and natural groundwater systems suggest a maximum U isotopic fractionation of
1.2‰ between aqueous U(VI) and reduced, particulate U(IV) (e.g., Bopp et al., 2009; Basu et al., 2014; Murphy et al., 2014; Stirling et al., 2015). However, due to diffusion processes in the sediment column, the maximum apparent fractionation factor likely to be observed is
0.6‰, or about half of the full fractionation factor (Clark and Johnson, 2008; Andersen et al., 2014). Additional constraints from U(VI)–U(IV) reduction in natural settings and experiments, largely without diffusion limitation, support a range of isotopic fractionation factors between U(VI) and U(IV) from 
0.3‰ to
1.0‰ (e.g., Rademacher et al., 2006; Bopp et al., 2009; Romaniello et al., 2013; Basu et al., 2014; Stirling et al., 2015; Noordmann et al., 2015b). In intermittently anoxic basins, such as Saanich Inlet in British Columbia and Landsort and Gotland Deeps in the Baltic Sea, U isotope fractionation between seawater and sediments yields variable fractionation factors (Holmden et al., 2015; Noordmann et al., 2015b), which may change significantly depending on bottom water renewal rates. In these basins, U removal from the water column is often negligible or below 10%. Thus,

the difference between permanent and intermittent anoxic conditions can significantly alter how U and its associated 238 U/235U composition is cycled through the water column and exported to sediments. It is therefore necessary to characterize additional basins of varying oxygen content to fully understand and quantify the mechanisms that drive changes in the d238U of the water column and underlying sediments before full application of the U isotope system as a paleoredox proxy can occur. Here we present paired trace metal concentrations of redox-sensitive metals (U, Mo, Fe, and Mn) and U isotope measurements (238U/235U) for the water columns and sediments of New Zealand fjords, ranging from fully oxic settings to intermittently anoxic/euxinic (H2S-bearing). Furthermore, we explore the link between intermittent anoxia and primary productivity, including the relationship with trace metal cycles. By constraining the U isotope response to intermittent anoxia in these basins, we aim to address the following outstanding research questions: (1) How do U, Mo, Fe, and Mn concentrations and U isotopes vary in the water columns and sediments across a range of oxygen conditions in New Zealand’s fjords? (2) Do other factors, such as primary productivity in the water column, affect U isotopic fractionation between seawater and underlying sediments? And (3) In combination with other studies of modern intermittently anoxic basins (Holmden et al., 2015; Noordmann et al., 2015b), can we derive estimates of U isotopic fractionation ranges between seawater and sediments for these settings? Finally, we discuss the significant implications of our results for reconstruction of past global ocean oxygen conditions. 2. BRIEF PRIMER ON 238U/235U AS OCEAN REDOX TRACER Uranium is the heaviest naturally occurring element with two primordial isotopes, 238U and 235U (natural abundances of 99.28% and 0.72%, respectively) (Dempster, 1935; Katz and Rabinowitch, 1951). Both isotopes have long half-lives of greater than 108 years (Nier, 1939). Until recently, it was believed that the global 238U/235U ratio was invariant at 137.88 (Naudet and Renson, 1975; Cowan and Adler, 1976). However, more recent studies have shown permil-level variability in 238U/235U in near surface environments (Stirling et al., 2007; Weyer et al., 2008) and have given rise to a new, higher accuracy value of 137.797 ± 0.005 (Tissot and Dauphas, 2015) for the average terrestrial composition. When converted to delta notation, defined as in (1) below, the ‘‘bulk silicate earth” composition yields a d238U of
0.29 ± 0.03‰ relative to the U isotope standard CRM 145 (Tissot and Dauphas, 2015). 
238 235   U = U sample d238 U ¼
1  103 ‰ ð1Þ 238 U =235 U standard In marine anoxic environments, U(VI)–U(IV) reduction has recently been shown to have significant isotopic fractionation at the permil level (Stirling et al., 2007; Weyer et al., 2008; Romaniello et al., 2013; Andersen et al., 2014; Holmden et al., 2015; Noordmann et al., 2015b).

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Montoya-Pino et al., 2010; Brennecka, Herrmann et al., 2011; Kendall et al., 2013; Asael et al., 2013; Murphy et al., 2014; Dahl et al., 2014; Noordmann et al., 2015a). However, some studies have shown varying water column U isotope behavior in anoxic basins that receive intermittent bottom water renewal. For example, in the Saanich Inlet, British Columbia and the Landsort and Gotland Deeps, Baltic Sea, replenishment of U during renewal events leads to only minor or no U isotopic shifts in the water column, presumably due to the introduction of ‘‘unfractionated” open-ocean U to the water column (Holmden et al., 2015; Noordmann et al., 2015b). Moreover, other processes acting in addition to the U(VI)–U (IV) exchange reaction, possibly governed by conventional mass-dependent fractionation and small differences in the zero-point vibrational energies between light and heavy isotopes, can lead to U isotopic fractionation in the opposite direction to that expected from nuclear field shift effects. This would give rise to the preferential enrichment of heavy isotopes and elevated d238U signatures in the aqueous phase. In particular, U adsorption onto organic matter (OM) and/or oxyhydroxides may preferentially remove light

This process is thought to occur on particle surfaces primarily below the sediment-water interface (SWI) (Anderson, 1987; Barnes and Cochran, 1990, 1993). The isotopic fractionation associated with reduction would result in more negative d238U values in pore waters compared with the sediment, which may then be transferred up the water column via diffusion processes (Clark and Johnson, 2008; Andersen et al., 2014). However, some modern water column profiles show decreasing d238U immediately below the chemocline (Noordmann et al., 2015b), suggesting the possibility that U reduction may occur on or within particles above the SWI as well. See Fig. 1 for schematic representation of expected U concentration and isotopic profiles under varying oxygen conditions. Uranium has a long oceanic residence time of ca. 400 kyr (Dunk et al., 2002) and is well-mixed in the open ocean, giving rise to globally homogeneous U concentrations and isotopic compositions. Thus, increases and decreases in the areas of suboxic and anoxic sinks will drive the global seawater d238U signature to more negative or positive values, respectively, making the 238U–235U system a potentially powerful global paleoredox tracer (e.g.,

(A) Oxic Conditions: Soluble, unreactive U(VI) in water column

(B) Suboxic Conditions: U(VI) reduced to insoluble U(IV) at or below sediment-water interface

(C) Anoxic/Euxinic Conditions: U(VI) reduced to insoluble U(IV) in water column or sediments

Sediment

0

≥5 mg/L 0

13.4 -0.8‰ nmol/L

-0.4‰

102-3 μM

[H2S]

δ238U

[Uaq]

[O2]

Water Column

0

O2

Little or no drawdown of U

0

≥5 mg/L 0

91

13.4 -0.8‰ nmol/L

-0.4‰

0

≥5 mg/L 0

Upward diffusion of porewater U?

13.4 -0.8‰ nmol/L

-0.4‰

Fe(III) → Fe(II) (zone of U reduction) H 2S

Fig. 1. Schematic representation of U concentration and d238U in the water column under varying oxygen conditions. Blue zone represents fully oxic conditions; orange zone represents the zone where oxygen is fully depleted and Fe(III) undergoes reduction to Fe(II), coincident with the zone of U(VI) reduction to U(IV); green zone represents euxinic conditions characterized by the presence of H2S. (A) Under fully oxic conditions, oxygen concentrations do not change down the water column. Aqueous U concentrations and d238U remain constant throughout the water column. No authigenic U is exported to sediment by redox-controlled processes. (B) Under suboxic conditions, the zone of U reduction sits near the sediment-water interface (SWI), causing minor drawdown of aqueous U from the lower water column. As reduction occurs at or below the SWI, sediment d238U shifts to more positive values while porewater d238U becomes more negative. (C) When anoxic conditions extend into the water column, aqueous U concentrations decrease measurably below the chemocline as U(VI) is reduced, converted to particulate form, and exported to underlying sediments. When export to sediment is controlled by redox-driven processes, the water column d238U should shift to more negative values as 238U is preferentially exported to particulate form. If anoxic/euxinic conditions persist in the water column, porewater U, with a more negative d238U than open ocean seawater, may diffuse back into the water column. This would produce an isotopically ‘‘light” water column, particularly below the chemocline. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

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235 U from seawater, acting to counter the signal of redoxdriven U isotope fractionation. One measurement of U isotope fractionation between seawater and plankton from Saanich Inlet, British Columbia indicates that seawater is 0.79 ± 0.17‰ heavier than plankton (Holmden et al., 2015). Additionally, experimental results show that U adsorption onto Mn-oxyhydroxides produces a measurable U isotopic fractionation of up to 0.2‰ where the light 235U isotope is preferentially adsorbed onto the solid phase, similar to a kinetic effect (Brennecka, Wasylenki et al., 2011). However, field observations of U adsorption onto clays suggest no measurable U isotope fractionation (Shiel et al., 2013). In settings where bottom water renewal, adsorption, and reduction all occur simultaneously, as in the New Zealand fjords, the U isotope systematics may be complicated by multiple competing processes, which act to obscure the d238U signature attributed to U reduction alone.

3. STUDY AREA 3.1. Fiordland National Park Fiordland National Park is located along the southwest coast of New Zealand’s South Island (Fig. 2). Fourteen glacially carved fjords extend inland from the coast, oriented east to west. At fjord entrances, bedrock sills rise to 150–20 m below present sea level (mbsl), which limits exchange between fjord waters and the open ocean (Stanton and Pickard, 1981). The bedrock in the fjord catchments is primarily comprised of the Median Batholith and Western Province orthogneisses, granites, and metasediments (Allibone et al., 2007, 2009). The most significant unit is the Western Fiordland Orthogneiss, which extends from Sutherland Sound in the north to Dusky Sound in the south (Fig. 2) (Allibone et al., 2009). The region intercepts the Southern Hemisphere westerly winds (SHWW), leading to extreme regional orographic precipitation (>6 m year
1) (Salinger and Mullan, 2001) as moistureladen air hits the Southern Alps mountain range east of the fjords. High precipitation causes significant runoff of terrigenous material into fjord basins, including organic and lithogenic material. Geochemical characterization of source rocks and material entering the fjords can be found in Allibone et al. (2007, 2009), Smith et al. (2010), and Hinojosa et al. (2014, 2015). 3.2. Fjord hydrography New Zealand fjord water columns are stratified, with a seaward-flowing low-salinity layer (1–15 m) sitting on top of incoming saline seawater (Stanton and Pickard, 1981; Stanton, 1984; Pickrill, 1987). This ‘‘conveyer belt” circulation pattern is restricted by fjord entrance sills, which force seawater to overtop bathymetric constraints to enter the fjords. Once seawater surges over sills, it sinks to the fjord basins, which reach up to 450 mbsl (Stanton and Pickard, 1981). If intervals between bottom water renewal are less frequent, the deep water in fjord basins can become oxygen-limited (Stanton and Pickard, 1981; Schu¨ller

et al., 2014). While little information on the frequency of bottom water renewal exists, a study in Milford Sound suggested a minimum of one renewal event per year based on biological respiration rates (Garner, 1964). Given previous evidence for the development of euxinia in some NZ fjord basins (Hinojosa et al., 2014), biological consumption of water column oxygen must be extremely rapid, and lateral water movement very restricted, to reach sulfidic conditions on sub-annual timescales. Intervals between bottom water renewal can also be inferred from temperature and salinity profiles. As suggested by Stanton and Pickard (1981), bottom waters are likely to be more dense and more saline after renewal events. After longer intervals of no renewal, vertical mixing with fresh surface waters will decrease both salinity and density at depth. Yet making this comparison requires multiple measurements of past temperature and salinity conditions within a single basin. Historical data of this type is limited in the NZ fjords. Comparison of absolute temperature and salinity values between different fjords also must take into account relative freshwater influx based on catchment sizes (Stanton and Pickard, 1981), as a higher freshwater input ratio (or a higher ratio of catchment area to total fjord area) will reduce salinity throughout the water column. 3.3. Sampling sites In this study, we present redox-sensitive trace metal and U isotopic profiles from water columns in four distinct fjord basins, ranging from fully oxic to anoxic/euxinic. Of the four, Long Sound basin in Preservation Inlet had the highest oxygen concentrations in the water column during sampling (>5 mg/l at the SWI). Long Sound is located in the southernmost fjord, which has an entrance sill depth of 30 mbsl (Stanton and Pickard, 1981). Preservation Inlet has a catchment area/fjord area (CA/FA) ratio of 6.0 (Stanton and Pickard, 1981), which is one of the lowest ratios for the region. This implies that freshwater input has a reduced impact on water column geochemistry and hydrography. The water sampling site is located 25 km from the open ocean and 11 km from the main river outflow. At this site, the SWI sits at 368 mbsl. The two suboxic water columns are Girlies’ Island, Dusky Sound and Deep Cove, Doubtful Sound. They are both inner fjord basins that reach oxygen levels of <2 mg/l at the SWI. At Girlies’ Island, the SWI is located at 169 mbsl, while at Deep Cove, the SWI is located at 93 mbsl. Dusky Sound has a CA/FA of 5.1 (Stanton and Pickard, 1981), which again implies a lower influence of freshwater influx on overall water column properties. The Girlies’ Island sampling site is located 42 km from the open ocean and 8.5 km from the main river outflow. Because of connectivity with Thompson Sound to the north, Doubtful Sound has a CA/FA of 8.0. However, the installation of a hydroelectric dam tailrace in the 1960 s led to increased freshwater outflow near the Deep Cove site, therefore the water column at this site has a thicker freshwater lens. The sampling site is located 40 km from the open ocean and 0.5 km from the tailrace outflow.

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Fig. 2. Location of sampling sites for this study. (A) New Zealand. B: Fiordland National Park with 1:2,000,000 geology map from GNS Science. Yellow dots represent sites of d238U surface sediment samples. (C–E): The fjords from which water column samples were collected (C: Doubtful Sound, D: Dusky Sound, E: Preservation Inlet). Purple stars represent sampling sites. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

Sportman’s Cove, Dusky Sound is an intermittently anoxic basin that reaches euxinic conditions (Hinojosa et al., 2014). It is located in a small inlet on the western side of Cooper Island, which lies in the middle of Dusky Sound. The entrance sill limiting exchange between Sportsman’s Cove and the main channel of Dusky Sound is 9 mbsl. The water column at this site is 35 m deep. Sportman’s Cove is located 31 km from the open ocean and 19.5 km from the main river outflow. In addition, the site receives direct terrestrial runoff from Cooper Island. Doubtful Sound, Dusky Sound, and Preservation Inlet contain multi-channel pathways that provide linkages between fjord waters and the open ocean. The deepest entrance sill that provides water to Deep Cove, Doubtful Sound is located at the mouth of Thompson Sound, located to the north of Doubtful Sound. A secondary sill at the connection of Thompson and Doubtful Sounds shoals to 90 mbsl, which restricts exchange between the two fjord systems. The mouth of Doubtful Sound has a deepest sill depth of 101 mbsl. Dusky Sound has an entrance sill depth

of 65 mbsl (Stanton and Pickard, 1981; Hinojosa et al., 2014), but water can enter the fjord at a deeper depth from Breaksea Sound, which has an entrance sill at 93 mbsl, via Acheron Passage. Preservation Inlet has a complex sill zone with several secondary sills that prevent water exchange with Long Sound and the open ocean. The outermost sill has a depth of 30 mbsl. Farther into the fjord, a constricted sill zone (<500 m across) at 34 mbsl limits water flow into Long Sound. Five surface sediment samples from the Caswell Sound watershed, Inner Nancy Sound, Sportsman’s Cove, Supper Cove, and Isthmus Sound were measured for their trace metal concentrations and U isotope compositions. See Fig. 2 for locations. 4. METHODS Sampling took place during April 2012 and July 2013 (austral autumn and winter, respectively) aboard the University of Otago’s R/V Polaris II. Water column and

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surface sediment samples were collected for trace metal elemental and U isotopic analysis. Additionally, particulate OM (POM) was collected from discrete depths in the Sportsman’s Cove and Girlies’ Island water columns to measure C concentrations in POM (CPOM) under oxic, suboxic, and anoxic conditions. 4.1. Sample collection 4.1.1. Water samples Water column samples were collected in the four studied fjord basins: Deep Cove (Doubtful Sound), Sportsman’s Cove (Dusky Sound), Girlies’ Island (Dusky Sound), and Long Sound (Preservation Inlet) (Fig. 2). At most sites, 2–20 liters of water were collected at discrete depths using Niskin bottles. At the Girlies’ Island site, in the top 40 m of the water column, water was pumped directly onboard using a weighted length of acid-cleaned polyethelene tubing, and then immediately filtered into sample bottles. Both the Niskin and pumping methods yielded similarly negligible contamination in elements of interest. Depths were selected based on Sea-Bird SBE19 CTD (conductivity-tem perature-depth) measurements to identify water column oxygen content, fluorescence peaks (measured with WETStar fluorometer, Wet Labs), salinity, and temperature. All sample processing onboard ship was conducted within a laminar flow workstation equipped with HEPA filtration. Half of each sample was filtered using acid-cleaned 0.45 lm cartridge filters to remove particulate matter, while the other half was stored unfiltered. All samples were acidified shipboard using high-purity 8 N HCl (purified using sub-boiling distillation and quartz stills) to a pH of 1.8, and then stored in acid-cleaned Nalgene high-density polyethylene (HDPE) bottles (for sample volumes <1 L) or Nalgene low-density polyethylene (LDPE) CubitainersTM (for sample volumes >1 L). 4.1.2. Water column particulates Additional seawater samples were filtered to collect particulate matter from discrete depths in the Sportsman’s Cove and Girlies’ Island water columns. Sampling in April 2012 occurred during anoxic conditions at Sportsman’s Cove and suboxic conditions at Girlies’ Island, while sampling in July 2013 occurred under oxic conditions in both basins. Approximately 5 L of water was collected from Niskin bottles and pumped through a vacuum manifold filtration system to trap particulates on pre-combusted Whatman glass GF/C filters (1.2 lm pore diameter). Increments of 500 ml of water were processed until flow became restricted, indicating the filter pores were clogged with particulate matter. The total volume filtered for each sample ranged from 500 ml to 5 L. Smaller volumes were filtered at depths of highest primary productivity, which is measured by the CTD fluorometer. Fluorescence peaks correspond to high levels of plankton-produced chlorophyll a (chl-a). Larger volumes were required for depths below the photic productivity zone. Once filters were clogged, several drops of 10% HCl were added to the filters to remove carbonate phases, which can alter the C isotope signal if present. Filters were dried in a low-temperature oven

(40 °C) overnight and stored in a cool, dry environment until returning to shore. 4.1.3. Sediment samples Surface sediment samples were collected, prepared, and analyzed as described in Hinojosa et al. (2014). To summarize, samples were collected using a Ponar grab sampler, or when possible, a box corer to preserve the SWI. After collection, samples were bagged and frozen. Plastic sampling tools were used to prevent contamination of samples. 4.2. Elemental concentration determination by ICP-MS 4.2.1. Water samples For the determination of dissolved metal concentrations, 0.5 ml aliquots of the filtered seawater samples were diluted by a factor of ten with 2% (v/v) HNO3 and measured for elemental concentrations with an Agilent 7500 cs/ce Quadrupole (Q-)ICP-MS at the Centre for Trace Element Analysis (CTEA), University of Otago. The instrument was operated in collision cell mode to minimize potential interferences of elements or molecules of overlapping mass to the elements of interest. For elements with post-dilution concentrations below 100 ppb, the Q-ICPMS was operated in pulse counting mode, while higher concentration elements were measured using analog mode. Sample measurements occurred in multiple cycles, with on-peak measurements for each element lasting 100 ms per cycle. The concentrations of 27 elements were quantified by ICP-MS and the results for the elements of interest are presented in Table 1. Of these, a suite of four redox-sensitive elements were the focus of the present study and comprise (detection limits in parentheses): Mn (2.5 ppb), Fe (5.0 ppb), Mo (0.10 ppb), and U (0.05 ppb). Concentrations were determined using an internal standard calibration regression. Three runs of MQ-water procedural blanks yielded no concentrations in the elements of interest above detection limits. Sample results were corrected for the factor-of-ten dilution to derive the original seawater concentration. Three samples were run in replicate to assess reproducibility. Note that some reported U concentrations were determined using techniques in isotope dilution and MC-ICP-MS, following the procedure outlined in Section 4.3. All reported uncertainties are 2 SD. 4.2.2. Sediment samples Frozen sediment samples were freeze-dried and homogenized. One hundred milligrams of dry sediment from each sample was digested with a 10:1 mixture of high-purity Teflon-distilled ca. 28 M HF and quartz-distilled ca. 15 M HNO3. Samples were then evaporated to dryness and refluxed in 6 M HCl and 7 M HNO3 for one day each to convert fluorides, formed during HF digestion, into readily soluble chlorides and nitrates. After refluxing, samples were evaporated to dryness and dissolved in 30 ml of 1.5 N HNO3. Aliquots of each sample were diluted with 2% (v/v) HNO3 and measured in duplicate for elemental concentration determination via Q-ICP-MS, using methods above. A basalt standard reference material, BCR-2, and two shale

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Table 1 Trace metal concentrations (Mo, U, Fe, and Mn), salinity, oxygen, d238U, and d234U profiles from the water columns of Long Sound (oxic), Girlies’ Island (suboxic), Deep Cove (suboxic), and Sportsman’s Cove (anoxic). Dashes indicate values below detection limits. Site

Depth (mbsl)

Salinity

Oxygen (lmol/L)

Fe (nmol/L)

Mn (nmol/L)

Mo (nmol/L)

U (nmol/L)

d238U (‰)

2 SE (‰)

d234U (‰)

2 SE (‰)

Long Sound

10 20 30 50 70 90 120 140 165

33.5 33.9 33.9 34.0 34.1 34.1 34.2 34.2 34.2

344 364 368 348 328 313 314 324 330

– – – – – – – – –

– – – – – – – – –

127 127 133 128 134 137 132 135 131


0.35
0.31
0.43
0.41
0.37
0.40
0.36
0.32
0.42
0.38 ± 0.09

0.06 0.06 0.07 0.07 0.06 0.06 0.06 0.06 0.06

147 152 155 158 149 151 149 150 158

5 5 6 6 5 6 6 6 6

Girlies’ Island (U concentrations of italicized samples from Q-ICP-MS)

10 20 30 40

34.3 34.4 34.5 34.6

356 393 383 369

– – – –

– – – –

133 139 128 129

12.318 12.201 12.343 12.496 12.456 12.526 11.771 11.499 11.910 Water column average: 12.031 11.8 11.9 11.9


0.35 – – –

0.06 – – –

148 – – –

5 – – –

51 70 90 110 125 140 150 157 165

34.7 34.8 34.9 35.0 35.0 35.0 35.0 35.0 35.0

364 322 268 232 201 164 151 128 106

– – – – – – – – – –

– – – – – – – – – –

129 128 122 136 133 127 129 136 122

– – – – – – – –
0.41
0.38 ± 0.08

– – – – – – – – 0.06

– – – – – – – – 155

– – – – – – – – 5

Deep Cove

10 30 50 65 75 85

33.8 34.2 34.4 34.9 34.9 35.0

435 392 359 251 198 93

– – – – – –

– – – – – –

132 132 125 127 128 127

11.9 11.5 11.4 11.9 11.7 11.5 11.5 11.6 11.905 Water column average: 12.073 12.000 11.892 12.465 11.924 11.373 Water column average:


0.42
0.45
0.34
0.37
0.38
0.38
0.39 ± 0.07

0.07 0.07 0.06 0.06 0.06 0.06

147 148 149 147 149 155

6 6 5 5 5 5

Sportsman’s Cove

10

34.1

361

96

48

127

11.897


0.41

0.07

145

6

20 24 28 32

34.5 34.7 34.8 34.8

71 3 0 0

901 1039 757 877

255 298 271 260

89 69 61 58

10.727 10.984 10.309 10.246 Water column average:


0.40
0.39
0.34
0.44
0.39 ± 0.08

0.07 0.06 0.07 0.06

145 155 151 151

8 5 6 5

80

29

15

MC (2 SE): 0.004 Q-ICP-MS: 0.6

Average 2 SD error

standard reference materials, USGS SGR-1b and SBC-1 (U.S. Geological Survey), were also measured. All elements of interest fell within the certified range for the three standards (Table 2). Additionally, total procedural blanks fell

below detection limits for all elements of interest. All reported uncertainties are 2 SD. Surface sediment concentrations of the trace elements Fe, Mn, U and Mo are normalized to Al concentration

– – – 5 129 146 132 10
3 – – – – – 0.07 0.05 0.05 0.05 0.05 0.08 – – – – –
0.27
0.26
0.64
0.14
0.32
0.32 – – 4.8E
05 1.1E
05 6.8E
06 3.1E
05 1.6E
05 2.4E
04 1.3E
03 1.7E
04 – – – 0.1 0.1 0.1 0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.2 0.5 1.2 0.2 0.3 8.7 4.3 0.8 6.8 1.9 0.2 2.5 4.1 9.7 1.1 2.9 61.2 22.6 1.6 262.4 37.3 3.0 5.6 5.6 5.7 7.4 5.8 1.9 0.6 5.4 6.9 3.1 10.0 Long Sound Girlies’ Island Deep Cove Caswell Sound watershed Inner Nancy Sound Sportsman’s Cove Supper Cove Isthmus Sound BCR-2 (basalt SRM) SGR-1b (shale SRM) SBC-1 (shale SRM) (NB: Shale errors are 1 SD Q-ICP-MS errors; no duplicates ran)

0.9 0.5 0.3 1.9 0.4 0.5 0.1 1.7 0.7 0.0 0.1

4.8 5.4 5.2 4.4 3.6 2.4 1.0 2.3 9.4 2.2 6.6

0.1 0.5 0.8 0.5 0.2 0.2 0.6 1.0 0.3 0.0 0.1

1.8E+00 2.0E+00 1.9E+00 1.2E+00 1.3E+00 2.6E+00 3.3E+00 8.8E
01 – – –

7.1E
02 6.4E
02 7.8E
02 8.8E
02 2.8E
02 1.8E
01 5.0E
01 1.2E
01 – – –

1753 853 665 771 461 201 73 322 1539 236 1076

267 154 62 219 105 52 98 76 125 4 17

6.4E
02 3.1E
02 2.4E
02 2.1E
02 1.6E
02 2.1E
02 2.4E
02 1.2E
02 – – –

3.5E
03 1.5E
03 6.3E
04 2.0E
03 9.5E
04 2.0E
03 8.2E
03 1.2E
03 – – –

1.6E
04 2.7E
04 5.7E
04 5.3E
05 1.7E
04 1.1E
02 1.4E
02 9.7E
05 – – –

1.7E
05 1.9E
05 3.5E
05 4.0E
06 7.4E
04 8.2E
04 8.0E
04 1.5E
05 – – –

3.9 3.1 2.3 3.3 1.9 7.9 12.4 5.7 1.6 4.7 5.3

6.0E
04 4.8E
04 3.6E
04 1.4E
04 2.8E
04 3.7E
03 1.6E
02 7.5E
04 – – –

d234U (‰) ±2 SE d238U (‰) ±2 SD U/Al ±2 SD U (ppm) ±2 SD Mo/Al ±2 SD Mo (ppm) ±2 SD Mn/Al ±2 SD Mn (ppm) ±2 SD Fe/Al ±2 SD Fe (%) ±2 SD Al (%) Site/Sample

Table 2 Trace metal concentrations (Mo, U, Fe, and Mn), Al concentrations, d238U, and d234U values, and molar ratios (Mo/Al, U/Al, Fe/Al, and Mn/Al) for fjord surface sediments.

– – – 6 5 5 5 4 5 – –

J.L. Hinojosa et al. / Geochimica et Cosmochimica Acta 179 (2016) 89–109 ±2 SE

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and reported as molar ratios. Aluminum is used as a tracer of detrital input because of its lithophilic nature, therefore normalizing to its concentration accounts for detrital Fe, Mn, U and Mo input into each basin (Tribovillard et al., 2006). In some cases, this normalization method can lead to spurious results (Van der Weijden, 2002). This possibility is reduced in settings where there is significant detrital input and where the coefficient of variation (CV) for Al sediment concentration (standard deviation divided by the mean) is significantly lower than that of the sediment trace element concentrations (Van der Weijden, 2002; Tribovillard et al., 2006). Due to high regional precipitation and proximity to river outflows, Fiordland meets the first criterion. Furthermore, the regional CV for Al is 0.36, which is generally lower than the regional CV for Fe of 0.36, Mn of 0.74, U of 0.86 and Mo of 2.39 (calculated from 24 surface sediment concentrations presented in Hinojosa et al. (2014)). Therefore, the use of the Al normalization method is justified in this region. 4.3. Particulate OM carbon isotopic analysis Dried filters of POM were loaded into silver capsules with small amounts of V2O5 to aid combustion for analysis. Carbon concentrations and isotopes were measured with an elemental analyzer connected to a continuous-flow isotope ratio mass spectrometer (CF-IRMS) at the University of Otago Isotrace Research Facility. 4.4. Uranium isotopic analysis by MC-ICP-MS Sample sizes corresponding to 50 ng of U were targeted for U isotopic analysis. This corresponded to a volume of 15–20 ml per seawater sample. Aliquots of 2–23 ml were required for the sediment digests, which were stored in 30 ml of 1.5 N HNO3. All chemical preparation for U isotopic analysis was performed in an ultra-clean, metal-free Class 10 (ISO 4) laboratory at the CTEA. All acid reagents used were of ultra-high purity and purified by sub-boiling distillation using quartz and/or Teflon stills, and all water used was Millipore Milli-QTM Element ultra-high purity (18.2 MX cm) water. For 238U/235U isotopic analysis, a mixed double-spike of isotopes 233U and 236U was added to each sample to correct for any isotopic mass fractionation that occurred during chemical extraction, if any, and instrumental analysis. The two spike isotopes, 233U and 236U have a near 1:1 ratio (236U/233U = 0.957111 ± 0.000054), while 235U and 238U are only present in trace amounts in the double-spike tracer (236U/235U = 1712.3; 236U/238U = 633.3) (Stirling et al., 2007). The 236U/233U ratio of the spike was calibrated against the uranyl nitrate standard reference material CRM 145, assuming a 238U/235UCRM145 of 137.837 ± 0.015 (Richter et al., 2010). A spike/sample 236U/235U ratio of 2.5 was targeted for all samples. After addition of the 233U–236U mixed tracer, the samples were refluxed overnight at 70 °C, evaporated to dryness to ensure sample-spike equilibration, then re-dissolved in 1.5 N HNO3. A two-step ion exchange chromatography protocol using TRU- and UTEVA-resin (Eichrom Technologies,

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U.S.A.) was employed to chemically extract and purify U from the sample matrix, following methods reported previously (Stirling et al., 2006, 2007; Amelin et al., 2010; Murphy et al., 2014). Purified samples were evaporated to dryness and dissolved in 0.25 N HCl + 0.05 N HF in preparation for MC-ICPMS analysis. Uranium isotope measurements were performed with a Nu Instruments Nu Plasma-HR MC-ICP-MS at the CTEA, following a modified method of Stirling et al. (2007). In brief, the isotopes of 238U, 236U, 235U, 234U, 233 U, and 232Th were measured simultaneously in Faraday cups operating with 1011 X resistors. Samples were introduced using a Nu Instruments DSN-100 desolvating nebulizer fitted with a Glass Expansion 50 ll PFA nebulizer, operating with a flow rate of 60–100 ll/min. CRM 145 was measured before and between sample runs to monitor instrument performance. CRM 145 provides a normalization standard by comparing the measured 238U/235U to the published value of 137.837 ± 0.011 (Richter et al., 2010). Additionally, procedural blanks were analyzed for U contamination. Blanks consistently fell below 5 pg of U, which represents a negligible contribution compared to the 50 ng sample size (contamination < 0.01% of total U). Therefore no external blank correction was applied to the U isotopic data. Within each sample run, online data reduction was performed to correct for the minor contributions of 238U and 235 U present in the double-spike solution. Also, a correction for instrumental mass fractionation was made by normalizing the measured 236U/233U ratio to the true value (Stirling et al., 2006) and the exponential mass fractionation law (Hart and Zindler, 1989; Habfast, 1998). The results of the U isotope measurements are reported in delta notation relative to the CRM 145 standard (New Brunswick Laboratory, USA) as described previously in Eq. (1). Sample d238U values were determined by normalizing to two runs of double-spiked CRM 145 that bracketed sample runs. All reported uncertainties for individual datapoints are 2 SE, representing the internal analytical error and the internal error of the bracketing CRM 145 standards, combined in quadrature. A long-term average d238U of 0.00 ± 0.10‰ (2 SD) was measured for CRM 145 using the same bracketing approach, by treating CRM 145 as an unknown sample. This is comparable to the average internal 2 SE uncertainty determined for individual measurements of ±0.08‰. In addition, a d238U =
0.32 ± 0.08‰ was measured for the United States Geological Survey (USGS) basalt standard reference material BCR-2, consistent with the previously reported range of d238U =
0.29 ± 0.07‰ (Weyer et al., 2008; Amelin et al., 2010; Condon et al., 2010; Murphy et al., 2014). In samples measured for U isotopic composition, U concentration was determined with high-precision by isotope dilution using the 238U/236U of the sample, corrected for both instrumental mass fractionation and the contribution of natural isotopes in the double spike, the calibrated 236 U concentration of the mixed spike, and the masses of both the sample and double spike. The method used to quantify the U concentration, isotope dilution using

97

MC-ICP-MS or external calibration using Q-ICP-MS, is reported in Table 2. Using the 234U ion beam intensities monitored in a Faraday cup, 234U/238U ratios were determined to assess relative contributions of U from detrital (d234U = 0‰) and authigenic (d234U = 149‰) (Andersen et al., 2010) sources. Results were corrected for the contribution of 234U, 235U and 238U present in the double spike then normalized to the composition of the CRM 145 U isotope standard, which has an expected 238U/234U of 18920 ± 2 (Andersen et al., 2010), in a similar manner as reported above for 238 U/235U values. The data are reported in delta notation relative to secular equilibrium (Sec. Eqm.) as follows: " 234 # ! U 238 U sample 234
1  103 ‰ d U¼ ð2Þ 234 U 238 U Sec:Eqm: Repeat measurements of CRM 145 yielded an average d234U of
38 ± 5‰ (2 SE). The basalt standard BCR-2 yielded a d234U = 0 ± 5‰, placing it within the range of Sec. Eqm., as expected. 5. RESULTS Water column trace metal concentrations, d238U, and d234U are presented in Table 1, and sediment trace metal concentrations, d238U, and d234U can be found in Table 2. 5.1. Trace metal concentrations in the water column and sediments 5.1.1. Iron and manganese In the oxic and suboxic water columns, Fe and Mn concentrations were below detection limits, indicating concentrations below ca. 90 nmol/L and 45 nmol/L, respectively. Below the chemocline in Sportsman’s Cove, concentrations of both elements rapidly rise to an order of magnitude above detection limits, reaching a maximum concentration of 1.039 nmol/L for Fe and 298 nmol/L for Mn at 24 mbsl. Concentrations then decrease slightly below this maximum, but stay above 750 nmol/for Fe and 250 nmol/L for Mn. Above the chemocline, low amounts of Fe and Mn were detected at 10 mbsl, reaching respective concentrations of 96 nmol/L and 48 nmol/L. In the sediments, molar ratios of Fe/Al indicate two basins with enriched Fe values relative to the previously published average of 1.43 ± 0.49 for surface sediment Fe/Al in Fiordland (Hinojosa et al., 2014). Specifically, Supper Cove and Sportsman’s Cove, both of which also have enriched concentrations of U and Mo, yield Fe/Al values of 3.3 and 2.6, respectively (Table 2). Iron enrichment at these sites provides evidence for Fe(II) deposition under euxinic, rather than suboxic/anoxic, conditions below the SWI and possibly in the water column. In particular, the presence of free H2S leads to the precipitation of Fe-sulfides (Berner, 1984), whereas under suboxic/anoxic conditions, Fe is released to the dissolved phase and depletion occurs in the underlying sediments. Molar ratios of Fe/Al for the other basins are considerably lower and range from ca. 0.9–2.0. Manganese displays reverse behavior,

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with Long Sound having the most Mn-enriched sediments (Mn/Al = 6.4  10
2), while Supper Cove and Sportsman’s Cove do not show significant Mn enrichment compared with other basins (Table 2). This indicates sediment deposition of particle reactive, oxidized Mn at oxic Long Sound, and reductive dissolution of Mn below the chemocline at suboxic Supper Cove and intermittently anoxic/euxinic Sportsman’s Cove.

4.4  10
3 (Hinojosa et al., 2014). In the oxic Long Sound basin, sediments yield a Mo/Al ratio of 1.6  10
4. In the suboxic Girlies’ Island and Deep Cove basins, Mo/Al ratios range from 2.7  10
4 to 5.7  10
4, which is 2–4 times higher than the oxic Long Sound sediments. At Sportsman’s Cove, the sediment Mo/Al ratio is an order of magnitude higher than the upper end of average oxic/suboxic range, with an Mo/Al of 1.1  10
2.

5.1.2. Molybdenum In the oxic and suboxic basins, Mo concentrations remain within a range of 120–140 nmol/L (Fig. 3A–C) throughout the water column, similar to the previously reported seawater average of 115 nmol/L (Siebert et al., 2003). Within each water column, Mo concentrations are all identical within their 1 SD error range. In contrast, anoxic Sportman’s Cove shows decreasing Mo concentrations near and below the chemocline (Fig. 3D). Here, the uppermost sample at 10 mbsl, within the oxic part of the water column, yields a Mo concentration of 127 nmol/L. This falls within the range measured at the other three oxic and suboxic sites. At 20 mbsl, near the chemocline, Mo concentration drops to 89 nmol/L. Below the chemocline, Mo concentrations continue to decline, reaching a minimum concentration of 58 nmol/L at 32 mbsl. This indicates a 55% decrease in dissolved Mo concentration from the upper oxic water column to the lower euxinic water column. Average surface sediment Mo/Al ratios in Fiordland range from zero (i.e., Mo below detection limits) to

5.1.3. Uranium Dissolved U has a consistent concentration of 12.1 nmol/L in near-surface waters (10 mbsl) in all four basins. This is slightly lower than the previously published open ocean average of 13.4 nmol/L (Ku et al., 1977; Dunk et al., 2002), but surface waters in the fjords are diluted by low U freshwater. Normalizing surface water U concentrations to a salinity of 35.00 to account for this dilution, where [U]norm = [U]  (35.00/measured salinity) (Ku et al., 1977; Owens et al., 2011), results in a higher adjusted U concentration average of 12.4 nmol/L. This salinitynormalized U concentration remains marginally lower than typical open ocean values. This may reflect minor U removal under suboxic conditions in the studied fjord locations, as this normalization method is only applicable in oxic settings, where U behaves as a conservative element and thus tracks salinity. In the oxic and suboxic water columns (Fig. 4A–C), U concentrations fall within a range of 11.4–12.3 nmol/L. In the deeper water columns of the oxic and suboxic sites of Long Sound, Girlies’ Island and Deep

Fig. 3. Molybdenum and oxygen concentrations in the oxic, suboxic, and anoxic water columns. Black circles represent surface sediment Mo/Al values for each basin. Dotted line represents the sediment-water interface. Note broken scale in A.

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Fig. 4. Uranium and oxygen concentrations in the oxic, suboxic, and anoxic water columns. A: Long Sound, B: Girlies’ Island, C: Deep Cove, D: Sportsman’s Cove. Black circles represent surface sediment U/Al values for each basin. Dotted line represents the sediment–water interface. Note broken scale in A.

Cove, U concentration appears to decrease slightly toward the SWI, below which U reduction is likely to occur (see Fig. 1). Additionally, at Long Sound and Girlies’ Island, the sample closest to the SWI shows an increase in U concentration, perhaps reflecting upward diffusion of U from pore waters below the SWI. The highest level of U removal from the water column occurs at the anoxic site, Sportsman’s Cove (Fig. 4.D), where below the anoxic interface at 20 mbsl, U concentrations decrease to a minimum of 10.2 nmol/L at 32 mbsl. This constitutes removal of 15% of the total dissolved U below the chemocline. The U/Al molar ratios of surface sediments at Long Sound (oxic), Girlies’ Island (suboxic), and Deep Cove (suboxic) range from 3.6 to 6.0  10
4 (see Table 2 for full values and errors). These values fall within the previously reported average U/Al range for Fiordland of zero (i.e., U below detection limits) to 1.34  10
3 (Hinojosa et al., 2014). The U/Al ratios at the oxic and suboxic sites are an order of magnitude lower than the sediment U/Al of 3.8  10
3 measured at anoxic Sportsman’s Cove. 5.2. d238U and d234U in the water column and sediments At the oxic to suboxic sites, d238U deviates little from the open ocean average of
0.42 ± 0.07‰ (Stirling et al., 2007; Weyer et al., 2008; Andersen et al., 2014, 2015; Tissot and Dauphas, 2015), giving rise to water column averages of


0.38 ± 0.09‰,
0.38 ± 0.08‰, and
0.39 ± 0.07‰ at Long Sound, Girlies’ Island, and Deep Cove, respectively (Fig. 5). Additionally, at the Sportsman’s Cove anoxic site, there is no measurable change in d238U below the chemocline at 20 mbsl, yielding a water column average of
0.39 ± 0.08‰. Thus, at all four sites of varying oxygen content, d238U remains within the range of reported seawater d238U values. Additionally, d234U values clustered around the seawater value of 147‰ (Andersen et al., 2010), resulting in a regional range of 151 ± 8‰ (Table 1) and providing an end-member value for authigenic U in surface sediments. Of the five measured surface sediment samples, three clustered within a d238U range of
0.28 ± 0.05‰, including a sediment sample from a riverbank upstream of Caswell Sound (Fig. 6). These d238U values fall within the previously reported range of
0.20‰ to
0.65‰ for granites (Weyer et al., 2008; Telus et al., 2012; Murphy et al., 2014), which, along with diorites and gneisses, form a significant part of the basement geology in the Fiordland region. Another method to determine detrital versus authigenic U contributions is the d234U of the total sediment. Detrital inputs from geological units older than ca. 600,000 years, as is the case for Fiordland basement rocks, should be in, or close to, radioactive secular equilibrium with respect to 238U and 234U, giving rise to a d234U of 0‰, whereas authigenic U derived from seawater and

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Fig. 5. Uranium concentrations (blue diamonds) and isotope ratios (d238U, black squares) in the oxic, suboxic, and anoxic water columns. The gray dotted line represents mean seawater d238U (
0.42‰) and its 2 SD uncertainty limits are denoted by the shaded gray region. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

Fig. 6. Uranium enrichment in sediments (U/Al) versus d238U in sediments. Blue dotted line represents the regional seawater d238U average (
0.38 ± 0.08‰), from which authigenic U in sediments have precipitated from. This process may be accompanied by isotopic shifts in d238U, the magnitude and direction of which are a function of the mechanism(s) controlling the system. Inner Nancy Sound, Caswell Sound watershed (terrestrial end-member), and Isthmus Sound show little authigenic U enrichment and cluster in the d238U range reported for granites (
0.2‰ to
0.65‰, (Weyer et al., 2008; Telus et al., 2012; Murphy et al., 2014)), which reflects the underlying geology of Fiordland. Sportsman’s Cove has a higher U/Al but d238U fractionates in the direction expected for U(VI) adsorption, not U(VI)–U(IV) reduction. Supper Cove also has a high U/Al but d238U fractionates in the direction expected for U(VI)–U(IV) reduction. All d238U error bars are 2 SE; all U/Al error bars are 2 SD. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

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deposited in modern to recent sediments should have a d234U signature close to the seawater value, which averages 151‰ in this region. Using these two end-member values, relative contributions of lithogenic and authigenic U can be estimated. Both Isthmus Sound and the Caswell Sound watershed have low d234U values of 10 ± 4‰ and 5 ± 6‰, respectively, indicating an almost exclusive contribution of detrital U to the total sediment (0–9% authigenic U). The d234U values thus agree well with the d238U signatures, as both are in the expected range for the regional lithology. The two sites characterized by water column anoxia/ euxinia, Supper Cove and Sportsman’s Cove, have enriched U/Al molar ratios as well as fractionated d238U. At both sites, d234U values fall very near or within the range of modern seawater, suggesting that 84–100% of U is authigenic and detrital contributions are minor (Fig. 7). Supper Cove sediments yield a higher d238U relative to seawater of
0.14 ± 0.05‰, while the d238U of Sportsman’s Cove sediments are lower than seawater at
0.64 ± 0.05‰. Although both anoxic sites display enriched U/Al ratios compared with the sediments in the oxic and suboxic basins, Supper Cove sediments are more enriched in U by a factor of 4 compared with Sportsman’s Cove (U/Al of 1.6  10
2 versus 3.6  10
3). The fifth site, Nancy Sound, has a d234U of 129 ± 5‰, which suggests a major contribution of authigenic U to the bulk sediments at this location (82–89% authigenic U). This site is not known to be anoxic, nor does the direction of d238U fractionation from seawater (DU(aq)-U(solid), defined as the d238U of aqueous seawater U minus d238U of particulate U =
0.2‰) suggest redox-driven fractionation. Rather, the d234U and d238U of the sample suggest high authigenic enrichment and adsorption processes, respectively, meaning the sediment U is primarily authigenic but enriched in 235U relative to seawater. It is possible

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that the depth of the chemocline is far below the SWI, which would reduce the apparent fractionation factor in surface sediments due to diffusion limitation (Clark and Johnson, 2008). However, more constraints on seawater and sediment conditions (e.g., oxygen concentrations, primary productivity estimates, accumulation rates, and oxygen penetration depth) would help to determine the leading U export process at this site. 5.3. Primary productivity in oxic, suboxic, and anoxic settings The Sportsman’s Cove and Girlies’ Island water columns were sampled during contrasting low-oxygen (anoxic or suboxic) and fully oxic conditions to investigate associated changes in primary productivity and particle export in relation to changing water column oxygen content (Figs. 8 and 9). In April 2012, Sportsman’s Cove reached total oxygen depletion from 20 mbsl to the SWI at 35 mbsl. During the same cruise, Girlies’ Island reached suboxic conditions (1.4 mg/L) near the SWI at 167 mbsl. In July 2013, when the two basins were sampled again, both water columns were fully oxygenated (>8 mg/L). Under anoxic conditions in Sportsman’s Cove, a chl-a peak (13.6 mg/m3) was measured below the chemocline at 24 mbsl. Concentrations of CPOM were also highest at this depth, reaching 389 lgC/L. CPOM in the oxic portion of the water column were also higher than concentrations measured in July 2013 by a factor of two or more. Carbon concentrations ranged from 101 lgC/L at 2 mbsl to 285 lgC/L directly above the chemocline (18 mbsl). Near the SWI, CPOM concentrations were 267 lgC/L. When sampled under oxic conditions during July 2013, chl-a and CPOM concentrations at Sportsman’s Cove were significantly lower, and variability down the water column

Fig. 7. d238U versus d234U in Fiordland sediments. Sediments with exclusively authigenic U should fall within the regional d234U range for seawater, shown by the black dotted line (average) and gray bar (2 SD error limits). Sediments at the two anoxic sites, Sportsman’s Cove and Supper Cove, are comprised almost exclusively of authigenic U, as indicated by their near seawater d234U values, but their corresponding d238U signatures fractionate in opposite directions away from the seawater value, and towards isotopically light and heavy compositions, respectively. Sediments of the terrestrial end-member site, Caswell Sound, have a d234U of 5 ± 6‰, in the range of secular equilibrium (0‰), and a d238U of
0.27 ± 0.07‰; both U isotope signatures for Caswell Sound are indicative of U sourced from granite rocks.

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Fig. 8. Comparison of carbon concentrations from particulate OM in Sportsman’s Cove. The top panel presents results from an anoxic interval in the water column (April 2012), whereas the bottom panel presents results from the same water column when fully oxic (July 2013). The gray horizontal line on the top panel represents the depth of total oxygen depletion.

Fig. 9. Comparison of carbon concentrations from particulate OM at Girlies’ Island. The top panel presents results from a suboxic interval in the water column (April 2012), whereas the bottom panel presents results from the same water column when fully oxic (July 2013).

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was also reduced. Chl-a displayed no apparent peak and had a water column average of 2 mg/m3, six times lower than the peak measured in the same water column under anoxic conditions. Concentrations of CPOM were also lower and fell within a restricted range of 45–95 lgC/L, with the highest concentrations in the near-surface sample (2 mbsl) and close to the SWI. Salinity profiles from each sampling interval indicate a freshwater layer of 5–10 m, with freshwater dilution continuing throughout the water column. In comparison, the Girlies’ Island water column showed similar vertical profiles of chl-a and CPOM concentrations under suboxic and fully oxic conditions (Fig. 9). In both cases, near-surface waters yielded peak chl-a values of 5.2 mg/m3 and 3.9 mg/m3, respectively. At 2 mbsl, CPOM had high concentrations during both oxygenation regimes. The structure of chl-a and CPOM distributions were similar during both sampling intervals, with high chl-a and CPOM concentrations near the surface, a mid-depth decrease, and a return to near-surface values close to the SWI. Surface salinities were lower during the July 2013 sampling cruise, which may have been a result of recent precipitation. In the suboxic water column, salinities remained below 35.0 until 70 mbsl, whereas that threshold was reached at 25 mbsl during the oxic sampling cruise. 6. DISCUSSION 6.1. Does trace metal behavior conform to expected redox systematics? 6.1.1. Trace metal concentrations In the oxic Long Sound and suboxic Girlies’ Island and Deep Cove water columns, U concentrations profiles behave semi-conservatively, as salinity-adjusted U concentrations remain close to the open ocean average value of 13.4 nmol/L, especially at near-surface and mid depths in the water column. At both locations, U concentrations decrease by 10% from 12.0 to 12.3 nmol/L in surface waters to 11.4–11.6 nmol/L deeper in the water column, although there is a slight increase in the near-SWI sample from oxic Long Sound and suboxic Girlies’ Island. This may be due to reductive dissolution of Mn- or Feoxyhydroxides in the sediment, benthic remineralization of OM, which would release adsorbed U back into the water column as dissolved U(VI) (Barnes and Cochran, 1993), or diffusion of U(VI) released from the underlying sediments back into the bottom waters of these fjords. The anoxic water column at Sportsman’s Cove experiences 15% U removal below the chemocline. Despite U(IV) being the thermodynamically favorable species under reducing conditions, permanently euxinic basins, such as the Black Sea, only export 40% of dissolved U to sediments (Anderson et al., 1989). This leaves 60% of U in the soluble U(VI) form, which means removal of U to sediments, even in highly reducing, closed-basin conditions, is not quantitative. Several factors could affect the magnitude of the U removal flux from the water column to the sediment. First, Sportsman’s Cove is intermittently anoxic/euxinic, rather than permanently anoxic, and the time interval between

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bottom water renewals will alter U concentrations in the water column. Each renewal event reintroduces oxygen and U at open ocean seawater concentrations, both of which become more depleted as the time since renewal lengthens through continuous removal of U to the particulate phase. To this end, sampling soon after a renewal event will result in a dampened water column signal of U removal with respect to both U concentration and isotopic composition, as isotopically fractionated U residing in the water column is intermixed with newly introduced unfractionated oxic seawater. However, evidence of authigenic U enrichment, and potentially isotopic fractionation, will remain in the underlying sediments if they are below the oxygen penetration depth (Cai and Sayles, 1996; McManus et al., 2005). Second, the dominant process leading to U enrichment in the underlying sediments may also affect the magnitude of the flux. In Sportsman’s Cove, U/Al is two orders of magnitude higher than in the oxic or suboxic basins (Fig. 6), confirming that authigenic U enrichment occurs in this basin. This is further supported by a d234U of 146 ± 5‰ for Sportsman’s Cove surface sediment, which indicates that seawater U (d234U = 149‰) is the primary source of sediment U at this site. While reduction of soluble U(VI) to particle-reactive U(IV) in the anoxic zone of the water column, or in the porewaters below the SWI, is one potential U removal mechanism, it is important to note that Sportsman’s Cove also has the most productive water column of the four studied fjord locations, leading to a large POM flux. Therefore, adsorption of U(VI) onto settling organic particles could also account for a large portion of U enrichment in sediments at the Sportsman’s Cove site, which has been suggested in previous studies of authigenic U accumulation (McManus et al., 2005; Andersen et al., 2014). For example, in the Black Sea, up to 54% of the U flux to the sediment may be due to adsorption onto sinking particles (Anderson et al., 1989). Additionally, experimental results suggest that significant U(VI) adsorption onto microbial biomass occurs under high biomass concentrations, characterized by accumulation rates that are faster than those associated with U(VI)–U(IV) reduction (Gorman-Lewis et al., 2005). A recent study measured significant U isotope fractionation associated with this process of biomass-U adsorption, where the biomass d238U was
0.79 ± 0.17‰ lower than that of seawater (Holmden et al., 2015). If bottom water renewal events occur too regularly to allow (a) reduction and significant depletion of U (VI) in the water column, or (b) diffusion of aqueous U(VI) into the lower water column following U(VI)–U(IV) reduction in porewaters near the SWI, then adsorption may prevail as the dominant process exporting U to sediments. The uranium 238U–235U isotope system will help to distinguish between these competing processes (discussed in Section 6.1.2). Seawater Mo concentrations do not decrease until full anoxic/euxinic conditions are met. In Sportsman’s Cove, 55% of dissolved Mo is exported to the underlying sediment below the chemocline. While this proportion of removal is higher than that of U, it is still not the quantitative removal seen in permanently euxinic basins

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(Helz et al., 1996). Again, the rate of bottom water renewal may play a significant role in the extent of Mo removal. Specifically, if renewal events occur before the quantitative drawdown of Mo, then water column concentrations will return to near open ocean seawater values. If anoxic conditions persist, the severity of Mo removal will continue, but will not be complete and reach 100% if the basin has regular Mo replenishment from renewal events. Behaving in the opposite manner from U and Mo, Fe and Mn species in the reduced form are considerably more soluble than the particle-reactive oxidized species. Iron and Mn concentrations are below detection limits in oxic and suboxic waters but measurable in the anoxic water column at Sportsman’s Cove. Across the 20 mbsl chemocline at Sportsman’s Cove, Fe and Mn concentrations rapidly rise by over an order of magnitude for both elements. Additionally, low levels of both elements can be detected at 10 mbsl. Upward diffusion of dissolved Fe and Mn from below the chemocline is the likely source of Fe and Mn at shallower depths. Yet as the dissolved elements are oxidized, they will form insoluble oxides, which are likely to scavenge elements such as Mo and act as sinking particulate shuttles that carry trace metals to the SWI (Algeo and Tribovillard, 2009; Tribovillard et al., 2012). Therefore, the presence of Fe and Mn in the water column at Sportsman’s Cove suggests active redox cycling near the chemocline, which is also a probable cause for Mo drawdown at this site. However, this process may not serve as an effective sink for U (Algeo and Tribovillard, 2009). 6.1.2. Constraints from uranium isotopes Uranium isotopic values for the oxic and suboxic waters of all four fjord study localities, including the upper water column at Sportsman’s Cove, fall within previously published ranges (
0.42 ± 0.07‰) for oxic seawater (Stirling et al., 2007; Weyer et al., 2008; Andersen et al., 2014, 2015; Tissot and Dauphas, 2015) (Fig. 5). Due to the well-mixed, conservative nature of uranium in the oceans, d238U for oxic seawater should be globally consistent. However, once oxygen is depleted and uranium is reduced, water column U concentrations would be expected to decrease and the d238U of the water column would be expected to become lower as 238U is preferentially exported to the reduced, particulate U(IV) phase. Yet this behavior is not observed in the anoxic water column of Sportsman’s Cove. Instead, d238U stays within the oxic seawater range below the chemocline, similar to the water column in seasonally anoxic Saanich Inlet in British Columbia (Holmden et al., 2015), or shifts slightly to higher d238U values (Fig. 5.D). In other permanently anoxic/euxinic marine basins, water column d238U shifts toward lower values as 238U is preferentially removed to the underlying sediments (Noordmann et al., 2015b). Other indicators of a reducing environment at Sportsman’s Cove, including U and Mo depletion, and Fe and Mn enrichment in the water column, as well as U, Mo, and Fe enrichment in the underlying sediments, are at odds with the apparent lack of redox-driven fractionation seen in the d238U signatures of this locality. A Rayleigh distillation model describing kinetic reactions in a closed system was used to evaluate the evolving

d238U signatures of Sportsman’s Cove seawater (U(VI) reactant) and sediments (U(IV) product) as U-uptake reactions proceed (Fig. 10). Because redox-driven reactions generally lead to unidirectional, kinetically driven isotope fractionations, this model, based on the approach of Mariotti et al. (1981), was selected rather than an equilibrium approach. The model assumes a finite pool of reactant U(VI) with no replenishment during the model run. In basins that undergo intermittent bottom water renewal, such as Sportsman’s Cove, the model is ‘‘reset” during each renewal event, meaning the progression of the reaction and d238U signatures of seawater and surface sediments is highly dependent on how frequently these events occur. Under these assumptions, the range of fractionation factors (a, defined as 238U/235Ureactant/238U/235Uproduct) that best fit the Sportsman’s Cove water column and sediment data is 1.0000–1.0003, which is consistent with increased export of 235U to the sediment. This observation matches the measured d238U of
0.64‰ in the underlying sediment, which produces an a of 1.0003 at this site. Given the combined evidence from water column and sediment data, the most likely a value at Sportsman’s cove is >1.0000, which excludes biologically mediated, redox-driven processes as the primary mechanism of U removal at this site. The lack of d238U fractionation at Sportsman’s Cove, or slight 235U depletion in the water column, could be influenced by one, or a combination, of the following factors: (a) the predominant process of U removal, (b) the rate and magnitude of that process, and/or (c) an isotopically distinct U input. If reduction of U were the primary U removal mechanism in the Sportsman’s Cove water column, perhaps isotopic fractionation does not occur until particles are below the SWI. Yet if this were the case, upward diffusion of isotopically ‘‘light” pore waters into the water column would impart a measurable shift in d238U over time (Fig. 1C), which is not seen, even in the deepest water column sample at Sportsman’s Cove. For this scenario to hold, the implication is that water column renewal occurs sufficiently frequently to outpace diffusion, otherwise fractionated U isotopic signatures would be observed in the water column, especially near the SWI. However, because bottom water renewal rates are estimated to be (inter)annual events and U reduction and diffusion processes likely occur at much faster timescales, particularly in bioturbated sediments (Zheng et al., 2002), this scenario is unlikely. Alternatively, other U removal processes, such as U(VI) adsorption, could be outpacing U(VI)–U(IV) reduction at Sportsman’s Cove, leading to a dampened net d238U isotopic signal. High productivity in the Sportsman’s Cove anoxic zone (Fig. 8) supports the possibility of increased adsorption, which would be expected to leave the d238U of dissolved U(VI) in the water column unchanged (Shiel et al., 2013) or heavier (Brennecka, Wasylenki et al., 2011; Holmden et al., 2015). To date, a maximum DU(aq)
U(solid) of 0.27‰ (Brennecka, Wasylenki et al., 2011) has been documented for U(VI) adsorption onto inorganic particles. For U adsorption onto OM, a DU(aq)
U(solid) of 0.79‰ has been observed (Holmden et al., 2015).

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Uranium isotope results from whole sediment digests reported in this study show further evidence of competing U export processes at the studied localities. Relative to the Dusky Sound seawater average d238U of ca.
0.39‰, the sediments at Supper Cove are enriched in heavy 238U (d238U =
0.14 ± 0.05‰) with a DU(aq)
U(solid) of
0.25‰, consistent with a redox-driven, nuclear-volumedependent fractionation mechanism. The fractionation factor, a, at Supper Cove is 0.9998, in line with ranges from other reducing environments (Rademacher et al., 2006; Weyer et al., 2008; Murphy et al., 2014). Uranium concentrations in sediments are also higher at this site, as expected for anoxic settings, although no water column data exist for Supper Cove to test the extent of oxygen depletion and d238U changes in the water column. In contrast, sediments at Sportsman’s Cove show d238U shifts to lighter values relative to the overlying water column (DU(aq)
U(solid) = 0.25‰), consistent with the direction of adsorption-induced U isotope fractionation (Brennecka, Wasylenki et al., 2011; Holmden et al., 2015). The composition of the particulate matter that may be drawing dissolved U from the water column is unknown, but the sediment d238U may give indicate whether U adsorption is onto oxyhydroxides or OM. The direction of d238U fractionation is consistent with adsorption onto both types of particulates (i.e., lower d238U in sediments than water). The magnitude of fractionation (0.25‰) aligns more closely with estimates for adsorption

Fig. 10. Sportsman’s Cove water column d238U (light blue) versus fraction of aqueous U remaining, calculated from starting measured concentration of 11.90 nmol/L at 10 m water depth. Dark blue lines represent Rayleigh fractionation model trends for a values ranging from 0.9988, corresponding to a
1.2‰ fractionation between seawater U and particulate U, to 1.0004. All a values <1.0000 correspond to expected shifts under a reducing scenario (i.e., product d238U is more positive than unreacted seawater d238U). All a values P1.0000 are consistent with adsorption-driven removal of U from the water column. Seawater data support an a range of 1.0000–1.0002, with the black dashed line representing the best fit regression. The top shaded gray area represents the uncertainty using a modified York regression. Below, the black dotted line represents the evolving d238U of particulate U as more U is removed from the water column with an a of 1.0003. The black line and lower gray bar represent the Sportsman’s Cove surface sediment d238U range of
0.64 ± 0.05‰. With an a of 1.0003, 20% of water column U would have to be removed to produce the average sediment d238U. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)

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onto Mn-oxyhydroxides (fractionation between aqueous and adsorbed U of up to 0.27‰) (Brennecka, Wasylenki et al., 2011), yet U is not as strongly affected by oxyhydroxides formation as other elements, such as Mo (Algeo and Tribovillard, 2009). An alternative mechanism could be U isotopic fractionation associated with U adsorption onto settling organic particles, which could have a DU(aq)
U(solid) of up to 0.79‰ (Holmden et al., 2015). Uranium has a strong affinity for OM (Klinkhammer and Palmer, 1991; McManus et al., 2005), which could lead to larger rates of adsorption at a high primary productivity site such as Sportsman’s Cove. While the latter process is more likely at this site, the measured DU(aq)
U(solid) of 0.25‰ suggests some influence of competing processes after the initial adsorption, as well as other reaction parameters (e.g., temperature, reaction rates, initial water column composition). As suggested in recent literature, adsorbed U may then undergo reduction, which would act to decrease the magnitude of DU(aq)
U(solid), or even drive it towards negative values (Holmden et al., 2015; Stirling et al., 2015). Thus, sediment d238U at Sportman’s Cove supports the hypothesized two-step reaction mechanism, whereby U is first adsorbed to OM and then reduced on particle surfaces (Holmden et al., 2015). Though unlikely, riverine input could potentially provide an input flux with variable d238U, which would be reflected in the sediments even if a redox-related fractionation factor were constant. Previous studies suggest a global average d238U of
0.24‰ for riverine input (Stirling et al., 2007; Tissot and Dauphas, 2015; Noordmann et al., 2015a). However, it has been suggested that U from smaller rivers may deviate from this average (Noordmann et al., 2015a). If riverine d238U were variable in Fiordland, water column d238U at fjord locations that are proximal to the coastline should be distinct from more distal sites. However, all four sampled water columns fall within statistically indistinct ranges, both with respect to each other and to seawater. Furthermore, surface water d238U would be more variable than the close range seen in the four basins. Therefore, river inputs with heterogeneous d238U values in Fiordland are unlikely. The d238U of the detrital component of sediments may also skew the whole sediment d238U, particularly in regions of variable bedrock or areas influenced by continental weathering. Holmden et al. (2015) found a particularly low d238U of
0.83‰ for the detrital end-member in Saanich Inlet, which they attribute to isotopic fractionation as a result of weathering processes. Yet in Fiordland, the two sediment sites with almost complete detrital U based on d234U end-member mixing (Isthmus Sound and Caswell Sound watershed) also yield d238U values in the range for granitoid bedrock. 6.2. Assessing the global application of the d238U paleoredox proxy Recent studies of the d238U systematics in modern anoxic ocean basins have shown mixed results, especially in basins that receive intermittent bottom-water renewal (Holmden et al., 2015; Noordmann et al., 2015b).

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In Saanich Inlet, a seasonally anoxic fjord in British Columbia, sediments show evidence of reductive U isotope fractionation consistent with the nuclear volume effect (Holmden et al., 2015). The fractionation factor between aqueous U(VI) and reduced sediment U(IV) (DU(aq)
U(solid)) is
0.62 ± 0.17‰ (2 SD), statistically identical the fractionation factor measured in permanently anoxic Black Sea sediments (DU(aq)
U(solid) =
0.61±0.12‰, (Andersen et al., 2014)). Yet in the Saanich Inlet water column, there is no measureable U drawdown below the chemocline, and d238U remains within the open ocean seawater d238U range throughout the entire water column (Holmden et al., 2015). The authors attribute this difference in apparent fractionation factor in waters and sediments to a two-step reaction mechanism. The first step involves the adsorption of U(VI) onto particles in the water column, with U isotope fractionation associated with preferential partitioning of 235U in the adsorbed species of up to 0.79 ± 0.17‰ (Holmden et al., 2015). The second step occurs when U(VI) is reduced to U(IV) with a negative fractionation factor, opposite to the first step (Holmden et al., 2015). Importantly, the rates of each of these steps will control the net signal seen in the water column and underlying sediments. Another two-step mechanism has been proposed independently for recent microbial experiments of U reduction as well (Stirling et al., 2015). In that study, the authors found that both U(VI) reduction and U co-precipitation with carbonate phases cause decreasing aqueous U concentrations. While U reduction is the dominant cause of phase change at the onset of reactions, co-precipitation becomes relatively more important as more U(VI) is converted to a solid. As solid-phase U concentrations increase, surface areas upon which aqueous U can adsorb also increase, thereby accelerating co-precipitation of U (Stirling et al., 2015). A study from the brackish deep waters of the Baltic Sea shows similarly complex U isotope fractionation patterns (Noordmann et al., 2015b). In the Gotland Deep and Landsort Deep basins, U concentrations show no decrease down the water column despite euxinic bottom waters. At Gotland Deep, d238U shifts to slightly lower values of d238U =
0.50 ± 0.07‰ at 150 msbl from a surface d238U maximum of
0.36 ± 0.07‰ (Noordmann et al., 2015b), which is consistent with redox-driven fractionation. However, U concentrations do not decrease down the water column, as they would under a reducing regime. Additionally, the negligible to minor U isotope fractionation in the Baltic Sea water column of <
0.17‰ is smaller than shifts seen in permanently euxinic basins such as the Black Sea, or the Kyllaren Fjord in Norway, where d238U shifts by
0.28‰ and
0.35‰, respectively, between oxic and anoxic waters (Romaniello, 2012; Noordmann et al., 2015). Similar to Saanich Inlet, a stronger signal of U reduction is seen in the d238U signature of the underlying sediments of the Baltic Sea basins. For example, in a short sediment core from the Landsort Deep basin, sediments reach a d238U maximum of
0.13‰ that is significantly heavier than modern seawater, yielding a DU(aq)
U(solid) of
0.28‰, which is identical to the observed fractionation between seawater and Supper Cove sediments. However, some intervals

within the core also yield d238U values close to seawater (d238U =
0.39‰) (Noordmann et al., 2015b), despite having elevated U concentrations, which could be linked to intervals of increased adsorption or regular renewal of bottom waters. The d238U isotopic investigations from Saanich Inlet (Holmden et al., 2015), the Baltic Sea (Noordmann et al., 2015b), and Sportsman’s Cove (this study) highlight the possibility of competing influences on U isotopic fractionation in anoxic settings. Before wide application of d238U variability as a paleoredox tracer, several factors must be constrained. First, the fractionation factors for all U uptake/release processes, and how they change under differing environmental conditions, should be determined from experimental and field-based studies. This includes adsorption onto OM and other particulates, as well as ‘‘reverse” reactions such as oxidation of reduced U(IV) to U(VI), as to date, most studies have focused exclusively on U(VI)– U(IV) reduction. To this end, a single experimental study shows d238U fractionation of up to 0.3‰ associated with oxidation of solid U(IV), resulting in a shift toward heavier d238U values in the dissolved U(VI) product (Wang et al., 2014), in agreement with theoretical expectations. Additionally, recent evidence has shown the importance of biotic versus abiotic U reduction pathways (Stylo et al., 2015; Stirling et al., 2015). Both Stirling et al. (2015) and Stylo et al. (2015) suggest that only the biotic reduction pathway can impart a measureable shift in d238U, whereas abiotic U reduction can occur with no significant isotopic fractionation. Thus, an understanding of reduction pathways in natural environments may provide evidence for the expected magnitude of U isotope shifts that could result. Second, when applied as tracers of ancient redox conditions, d238U records should consider the relative effects of changing U input from terrestrial sources, U-OM adsorption processes, and local anoxic signatures. The first can be assessed through indicators of relative terrestrial and detrital input, e.g., concentrations of lithophilic minerals or terrestrial biomarkers. To account for potential complications of the d238U signal by OM adsorption, d238U records should be paired with other redox indicators, such as redox-sensitive element concentrations or other redoxsensitive isotope systems like d98Mo, as has been done in several d238U studies aimed at reconstructing ancient ocean redox conditions (e.g., Asael et al., 2013; Dahl et al., 2014; Kendall et al., 2015). Finally, to confirm that anoxic signatures at a single site are a result of global redox changes rather than local ones, studies that pair records from multiple sites will be the most robust. Evidence from this and recent studies of modern anoxic basins (Holmden et al., 2015; Noordmann et al., 2015b) highlight the potentially conflicting results when single basin dynamics are extrapolated to a global signature. 7. CONCLUSIONS The results presented here from oxic, suboxic, and intermittently anoxic fjords in New Zealand reinforce the need for improved constraints on the rates and magnitudes of reactions that drive changes in d238U in seawater and

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sediments. While the systematics of redox-sensitive trace metal enrichments are well understood, the stable isotope systems of most of these metals and their emerging application as local and global paleoredox tracers requires further exploration, based on a combination of experimental, theoretical, and field-based studies. From the present study, we draw the following conclusions in the context of the 238 U–235U isotope system. In low oxygen basins in New Zealand’s fjords, U and Mo are enriched in anoxic surface sediments, while Fe and Mn are depleted, consistent with redox-driven phase changes. Surface sediment redox-sensitive trace element concentrations from low-oxygen basins provide integrated evidence of past marine hypoxia, although cycling of trace metals into and out of sediment during bottom water renewal events is not yet fully constrained, which may affect both element concentrations and isotopic signatures. The 238U–235U isotope system displays conflicting signals in New Zealand fjord waters and sediments. Similar to the intermittently anoxic water columns in Saanich Inlet and the Baltic Sea, d238U shows little variability across the chemocline in the Sportsman’s Cove water column. This may be linked to (a) the intermittent nature of anoxia in this fjord due to the (inter)annual renewal of oxic waters (Stanton and Pickard, 1981; Pickrill, 1987) and the time since the last oxygenation event, (b) the character of the primary U removal process (e.g., U(VI)–U(IV) reduction versus U(VI) adsorption), and the sites where U removal occurs (e.g. in the anoxic/euxinic water column or below the SWI) or (c) a combination of these factors, and others. Two basins, Sportsman’s Cove and Supper Cove, both show systematic variations in trace metal concentrations indicative of reducing conditions, but the d238U results show opposite senses of fractionation from ‘‘unfractionated” seawater. Sportsman’s Cove yielded a DU(aq)
U(solid) of 0.25‰, consistent with adsorption-driven processes, while the DU(aq)
U(solid) at Supper Cove was
0.25‰, consistent with redox-driven processes. The competing influence of these processes and variability of other reaction parameters may lead to a wide range of apparent fractionation factors between seawater and sediments, especially in regions of high primary productivity where U-OM adsorption is likely. The d238U proxy for past global anoxia is a potentially powerful new tool to reconstruct and better understand ancient ocean chemistry. However, the present study reinforces the need for a more thorough understanding of the U isotope system, including improved constraints on the exact mechanisms controlling U isotope fractionation. As these various factors become more fully understood, the application of d238U will provide an independent method for reconstructing past ocean anoxia for comparison with other existing systems that are also becoming increasingly parameterized, thereby strengthening our understanding of ocean chemistry and links to the global carbon cycle and climate. ACKNOWLEDGEMENTS Funding for this research was provided by a University of Otago research grant (CHS and CMM) and a Royal Society of

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New Zealand Marsden Grant to CHS (UOO1314) and Fast-Start Marsden Grant to CMM (UOO1118). J.L.H. gratefully acknowledges support from a U.S. Department of State Fulbright grant. Many thanks to two anonymous reviewers and associate editor Dr. Shaw for thorough comments that greatly improved this work. We thank David Barr, Robert Van Hale, and Dianne Clark (Department of Chemistry, University of Otago) for laboratory and data analysis support. We are grateful to Candace Martin for use of her clean lab facility at the University of Otago for some aspects of sample preparation.

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