Deep-Sea Research I 112 (2016) 68–78

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Deep-Sea Research I

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Ocean productivity south of Australia during spring and summer

Michael L. Bender a,n, Bronte Tilbrook b, Nicolas Cassar c, Bror Jonsson a, Alain Poisson d,Thomas W. Trull b

a Department of Geosciences, Princeton University, Princeton, NJ 08544, USAb CSIRO Oceans and Atmosphere and Antarctic Climate and Ecosystems Cooperative Research Centre, Hobart 7001, Australiac Division of Earth and Ocean Sciences, Nicholas School of the Environment, Duke University, Durham, NC 27708, USAd Laboratoire IMAGES, Université de Perpignan Via Domita, 52 avenue Paul Alduy, 66860 Perpignan Cedex, Via Domitia, France

a r t i c l e i n f o

Article history:Received 24 November 2015Received in revised form23 February 2016Accepted 25 February 2016Available online 30 March 2016

x.doi.org/10.1016/j.dsr.2016.02.01837/& 2016 Elsevier Ltd. All rights reserved.

esponding author.

a b s t r a c t

We estimated mixed layer gross and net community production on a total of 20 crossings in the Aus-tralian sector of the Southern Ocean during the summer half-years (October–March) of 2007–2010. Theseestimates were calculated from measurements of O2/Ar ratios and triple isotope compositions of O2 in�250 seawater samples collected underway. For comparison purposes, we also measured the seasonaldrawdown of mixed layer −NO3 and SiO2 concentrations during 2006–2007 and 2007–2008. Across allsamples, average values of gross and net O2 production (measured by O2/Ar and O2 isotopes), were about86790 and 18717 mmol O2 m�2 day�1, respectively. Gross production was highest at the SubtropicalFront (up to �230 mmol O2 m�2 day�1), and decreased southward (to �10 near the southern boundaryof the Antarctic Circumpolar Current). In contrast, net community production showed little meridionalvariation. Net and gross O2 production increased throughout the spring-to-fall period, although mostSiO2 drawdown occurred in December. Consistent with satellite chlorophyll estimates, we saw no evi-dence for an intense spring bloom (e.g. as has been observed in the North Atlantic). Volumetric net andgross O2 production in the mixed layer, normalized to chlorophyll, increased (with considerable scatter)with average irradiance in the mixed layer. These relationships provide a basis for estimating productionfrom Argo float data and properties observed by satellite.

& 2016 Elsevier Ltd. All rights reserved.

1. Introduction

The Southern Ocean has long been identified as the largestHigh Nutrient Low Chlorophyll (HNLC) region in the global ocean.Variations in its status are important to nutrient delivery to otherocean basins (Sarmiento et al., 2004). They are also relevant toglobal carbon sequestration by the biological pump on decadal toat least interglacial timescales (Matear and Hirst, 1999; Sarmientoand Orr, 1991; Sigman and Boyle, 2000). The High Nutrient ap-pellation is an over-simplification. Phosphate and nitrate remainabundant across the entire Southern Ocean throughout the year.However silica, an essential nutrient for diatoms, which are thedominant large phytoplankton group, becomes depleted in sum-mertime waters between the Subantarctic Front and the PolarFront (Nelson et al., 2001; Treguer and Jacques, 1992; Trull et al.,2001a). Moreover, iron and light are actually the key limiting“nutrients”. Lack of iron, an essential element in enzymes, is nowknown to limit primary production essentially everywhere in theSouthern Ocean except near localized sources such as islands,

plateaus, and boundary currents (Boyd and Ellwood, 2010). Lowlight availability driven by both deep surface mixed layers andcloudiness also limits photosynthesis, including physiological in-teractions between Fe scarcity and the efficiency of light utilization(Hiscock et al., 2008; Mitchell et al., 1991; Strzepek et al., 2012;Sunda and Huntsman, 1997).

The Low Chlorophyll status is often linked to these limitationson primary production, but that is also an over-simplification,because loss by grazing and export is at least as strong a control onphytoplankton biomass accumulation (Cullen, 1995; Frost, 1991;Morel et al., 1991; Smith and Lancelot, 2004). Indeed, despite theLC status, Southern Ocean fluxes of organic carbon to the deepocean in sinking particles are close to the global median (Honjoet al., 2008, 2000; Trull et al., 2001c). The efficiency of this grazingand export may be modulated as mixed layer [SiO2] varies sea-sonally and meridionally.

Improving the understanding of the controls on the HNLCstatus of the Southern Ocean requires measurements of upperocean carbon fluxes. In this context, we present a large set ofpaired estimates of gross primary production (GPP) and netcommunity production (NCP) derived from gas tracers. The ob-servations cover the north-south extent of Southern Ocean waters

M.L. Bender et al. / Deep-Sea Research I 112 (2016) 68–78 69

and the late spring to early autumn half-year that dominates an-nual production. As outlined in the Methods section, GPP esti-mates are based on the isotopic composition of dissolved oxygen,and the NCP estimates are based on the excess of dissolved oxygenconcentrations relative to biologically inert argon concentrations.Conversion of these state variables to the GPP and NCP rate vari-ables is achieved by calculation of gas exchange rates with theatmosphere. Both terms reflect productivities averaged over thegas exchange time of the mixed layer (¼mixed layer depth/gastransfer velocity) or �2 weeks.

The results are based on discrete samples collected from theunderway clean seawater supply intake onboard the French Ant-arctic station resupply ship, Astrolabe. Astrolabe made threeround-trip crossings of the Southern Ocean each year betweenHobart, Tasmania, Australia and Dumont d’Urville, Adelie Land,Antarctica. These crossings made it possible to characterize GPPand NCP six times each season. We can thus evaluate spatial,seasonal, and interannual variations in the rate terms. Using sup-plementary surface seawater nutrient analyses, Argo float data,and satellite observations, we then can also explore relationshipswith surface phytoplankton biomass, light levels, and mixed layerdepths. Specifically, we focus on three questions:

1. Is there a spring phytoplankton bloom seen in primary pro-duction (as opposed to biomass accumulation) in the Australiansector of the Southern Ocean, and if so, what are itscharacteristics?

2. How does productivity vary with latitude in the Australiansector, and what factors mediate this relationship?

3. Is there a simple relationship between NCP and GPP and au-tonomously sensed properties, such as chlorophyll (Chl), sea-surface photosynthetically available radiation (PAR), and mixedlayer depth (MLD)? If so, can it be used, in conjunction withancillary data, to scale local estimates of NCP and GPP tobroader regions?

As expanded on in the Discussion section, our study regionsouth of Australia has the advantage of much background in-formation. Material includes (a) understanding of circulation fromsix repeats of the WOCE/CLIVAR SR3 hydrographic section fromTasmania to Antarctica (Rintoul and Trull, 2001; Sokolov andRintoul, 2002; Sokolov and Rintoul, 2009; Trull et al., 2001a),(b) special volumes from two month-long process studies of traceelement (Fe, Mn, etc.) and macro-nutrient (N, P, Si) biogeochem-istry and microbial tropho-dynamics in the SAZ (Bowie et al.,2011a, 2011b, 2009; Cassar et al., 2011, 2015; Sedwick et al., 2008,1999; Trull et al., 2001b), and (c) an ongoing program, as part ofAustralia’s Integrated Marine Observing System (www.imos.org.au), of air to sea and surface to deep ocean carbon fluxes at theSouthern Ocean Time Series site (Rigual-Hernández et al., 2015;Shadwick et al., 2015b; Trull et al., 2001c; Weeding and Trull,2014). These efforts provide important context for ourobservations.

2. Experimental methods

2.1. Principles of the dissolved O2 tracer methods for GPP and NCP

The O2/Ar ratio of dissolved gases in the mixed layer, and theδ17O and δ18O of O2, can be used to constrain mixed layer NCP andGPP, respectively (Castro-Morales et al., 2013; Luz and Barkan,2000; Prokopenko et al., 2011; Reuer et al., 2007). The O2/Ar ratioallows physical supersaturation (reflected in both O2 and Ar con-centrations) to be removed from the total O2 supersaturation. O2/Ar measurements thus quantify that part of O2 supersaturation

derived from photosynthesis in excess of respiration. Multiplyingbiological O2 supersaturation by the gas transfer velocity and sa-turation O2 concentration gives the sea-air flux of biological O2

(termed “net O2 bioflux” or just “net bioflux”). At steady state, andneglecting vertical mixing (see below), this flux must be sustainedby photosynthesis in excess of respiration, and hence provides anestimate of NCP.

The estimation of GPP is based on recognition that there aretwo sources of dissolved O2 in the mixed layer, atmospheric O2,and O2 produced in situ by photosynthesis. Measuring both δ17Oand δ18O of the dissolved O2 constrains the fraction deriving fromeach of these sources. The tracer quality of the two δ terms comesfrom the fact that O isotopes are mass-dependently fractionatedby photosynthesis and respiration (for which the change in δ17Ois�0.5 times the change in δ18O). However, photochemical ex-change in the stratosphere leads to non-mass-dependent fractio-nation, which in turn causes the global atmosphere to be depletedin 17O relative to 18O (Luz and Barkan, 2000; Luz et al., 1999). Therelative proportions of atmospheric and photosynthetic O2 in thedissolved O2 pool can be computed from the magnitude of the 17Odeficiency. Then, from the proportion of photosynthetic O2, the O2

concentration, and the gas transfer velocity, one can calculate thesea-air flux of photosynthetic O2 (“gross bioflux”). At steady-state,and again neglecting vertical mixing, this term provides an esti-mate of O2 gross primary production.

In the text that follows, we generally use the terms NCP andGPP to represent true values of production. We generally use theterms “net O2 bioflux” and “gross O2 bioflux” to represent the sea-air fluxes, calculated from observations, that approximate true netand gross production.

The estimates of NCP and GPP from the O2 measurements relyon the assumptions of biological steady-state (constant NCP andGPP) and negligible mixing across the base of the surface mixedlayer, as well as the parameterizations of air-sea gas exchange (asdetailed below). Probable biases in net O2 bioflux estimates of NCPhave been calculated in a model context (Jonsson et al., 2013). Onaverage, when O2 is biologically supersaturated in the mixed layer,net O2 bioflux underestimates NCP by about 10%. Additional ran-dom errors are much larger, probably about 740%, because ofnon-steady-state conditions, mixing across the base of the mixedlayer, and analytical uncertainties. Experimental studies also showthat net and gross bioflux can be influenced by upwelling, verticalmixing, and entrainment (e. g., Castro-Morales et al., 2013). Weminimize these issues by excluding from our analysis samples thatare undersaturated in O2 (which signifies upwelling). In any case,it is means and trends in our data that are robust, rather thanindividual excursions.

Nicholson et al. (2014) compared gross bioflux with GPP assimulated in a model context. In their simulations, gross biofluxunderestimated GPP in spring. This discrepancy was associatedwith a modeled spring bloom that is not observed in our data. Insummer and autumn, modeled gross bioflux was 2–3 times higherthan simulated GPP in the mixed layer. This error derived largelyfrom the entrainment and mixing of deeper euphotic zone watersinto the mixed layer, a phenomenon unlikely to be important inour study area. In our data set, MLD was440 m at 88% of XBTstations, and430 m at 80% of stations. According to Westwoodet al. (2011), most primary production in the Australian sectoroccurs above 30 m depth, and about 90% above 40 m depth.Therefore it is improbable that our data are seriously aliased byentrainment and mixing of deeper photosynthetic O2 into themixed layer. There are additional errors associated with lateraladvection and non steady-state conditions (Nicholson et al., 2014).Again, we need to focus on average values and trends rather thanindividual samples.

Fig. 1. Map of the study region showing the location of each water sample.

M.L. Bender et al. / Deep-Sea Research I 112 (2016) 68–7870

2.2. Dissolved O2 sample collection and analysis

The water samples (Fig. 1) were collected from the underwayclean seawater supply, which has its intake at about 3 m depth onthe resupply ship Astrolabe. Water samples were collected intopre-evacuated and HgCl2 pre-poisoned 500 ml bottles withLeuwers-Happert O-ring valves, as described by Emerson et al.(1995) and Reuer et al. (2007). The gases equilibrated passivelybetween the water and headspace before analysis in Princeton,and most of the water was removed by aspiration. Non-con-densable gases (essentially O2, N2, and Ar) were frozen onto mo-lecular sieve and then passed through a gas chromatographiccolumn to separate O2 and Ar from N2. The O2 and Ar were frozeninto a cryo-cooler at �15 K, warmed, and admitted to the FinniganMAT 252 mass spectrometer as described by Reuer et al. (2007).18O16O, 17O16O, and 16O2 were measured by simultaneous collec-tion for 3 cycles with 24 blocks per cycle. 17Δ of O2 is defined bythis equation:

Δ = ( × δ + ) − × ( × δ + ) ( )ln 0.001 O 1 0.516 ln 0.001 O 1 117 17 18

The uncorrected 17Δ of O2 was computed, using Eq. (1), fromthe raw data, based only on the observed ion current ratios. Anumber of factors can affect the raw values of δ17O and δ18O. Weevaluated these factors by analyzing the variability of between 76and 311 air standards measured during each of 5 analysis periods.In total, these extended from February 14, 2006, to January 1, 2011.

We rejected all samples where the mass 28 peak was greater than0.3% of the mass 32 peak, because high N2 leads to large errors inδ17O. We observed no dependence of δ17O and δ18O on the ratio of28–32 ion currents for the remaining samples, or on the initialpressure imbalance between sample and reference sides (data notshown). We did not correct for the zero enrichment, which ef-fectively cancels out when seawater is compared with waterequilibrated in the laboratory.

The 17Δraw values were corrected for differential gas depletionbetween sample and reference sides during analysis (Stanley et al.,2010), according to the formula:

( )( )

= −

– − ( )

− −

− −

Imbalance term 1 V /V sample

1 V /V 2

32sample end 32 sample start

32reference end 32reference start

The imbalance term registers the change in the ion currentimbalance between sample and reference during 1 block. “V” is thevoltage, and the subscripts “end” and “start” refer to the voltagesread at the end of the first block by the interfering masses programand that read at the start of the first block by the standard Fin-nigan ion current acquisition software from which δ values arecalculated.

17Δ data for a large number of equilibrated lab water samplesanalyzed in this way were higher than the equilibrium value(discussed below) by about 0.005‰. We therefore subtracted theaverage difference, calculated for each of our 5 analysis periods,from the corrected 17Δ value of O2. Knowing 17Δ and δ18O, wethen calculated δ17O from Eq. (1). Because of numerical issues, thisapproach is marginally more accurate than calculating 17Δ aftercorrecting the measured δ17O and δ18O values independently. Thechoice of the coefficient (in Eq. (1)) is immaterial to the result: wecalculate the same δ17O value if we use a coefficient of 0.518. Wedo not use 17Δ for any other purpose in the analysis of our results.

δ18O of O2 is fractionated by 0.1–0.3‰ during the process ofsample preparation. A correction to the measured δ18O value ismade which does not affect 17Δ. O2/Ar ratios were normalized toair and a correction was applied for the residual gas in the watersample after equilibration with the headspace. Uncertainties are73‰ for the O2/Ar ratio, 70.1‰ for δ18O of O2, and 70.008‰ for17Δ of O2.

2.3. Calculation of net and gross production from the O2 tracers

From δ17O and δ18O calculated as described above, the ratio ofG/k[O2]eq is calculated from Prokopenko et al. (2011, equation (7)),transformed to δ notation as described in their paper:

⎛⎝⎜

⎞⎠⎟

⎛⎝⎜

⎞⎠⎟

⎛⎝⎜

⎞⎠⎟

⎛⎝⎜

⎞⎠⎟

λ

λ=

− − −

− − −( )

δ

δ

δ

δ

δ

δ

δ

δ

+

+

+

+

+

+

+

+

−

−

−

−

−

−

−

−

GkO

1 1

1 13

O

O

O

O

O

O

O

O

2

10 1

10 1

10 1

10 1

10 1

10 1

10 1

10 1

eq

dis

eq

dis

p

dis

p

dis

3 17

3 17

3 18

3 18

3 17

3 17

3 18

3 18

G is gross O2 production, k is the gas transfer velocity(m day�1), and O2 is the saturation O2 concentration. The sub-scripts “p”, “eq”, and “dis” refer to photosynthetic O2, equilibriumO2, and dissolved O2, respectively. λ¼0.518. As shown in thisequation, G/k[O2]eq is a function of the equilibrium differencebetween the δ17O and δ18O of dissolved O2 and that of air O2, theδ17O and δ18O of dissolved O2 (which we measure), and the δ17Oand δ18O of photosynthetic O2. Air O2 is the reference.

δ18Op is set at the value for seawater relative to air,�23.320‰,essentially indistinguishable from the value of �23.324‰ re-ported by Barkan and Luz (2011). δ17O is then set to �11.856‰.With this assignment, the 17Δ of seawater and photosynthetic O2

with respect to atmospheric O2 is þ0.249‰. These settings invokethe assumption that the δ17O and δ18O of photosynthetic O2 is

M.L. Bender et al. / Deep-Sea Research I 112 (2016) 68–78 71

identical to that of the substrate water. We note the discovery thatphotosynthetic O2 is fractionated inside cells by respiration beforeit becomes part of the ambient O2 pool (Luz and Barkan, 2011).However, as long as the adopted parameters give a 17Δ value ofþ0.249‰ for seawater with respect to air, calculated values of G/k[O2]eq will be insensitive to the exact values selected for the iso-topic composition of photosynthetic O2.

Finally we need to adopt values for δ17Oeq and δ18Oeq. δ18Oeq

was accurately measured by as a function of temperature (Bensonand Krause, 1980), and we use their temperature-dependent va-lues. We adopt the 17Δ value for dissolved O2 in equilibrium withair, þ0.008‰, measured by Reuer et al. (2007) and subsequentlyby Stanley et al. (2010).

Gas transfer velocities between the ocean and atmospherewere calculated as weighted-average values for the 60 days beforecollection of the sample, using wind speeds from the NOAA/Sea-winds blended product (Chin et al., 1998), and the gas transferparameterization of Sweeney et al. (2007). Gas transfer velocitiescloser to days of collection were weighted more heavily (Benderet al., 2011; Reuer et al., 2007). The sea-air biological O2 flux (“netO2 bioflux”) was then simply calculated from the product of thebiological O2 supersaturation, O2 solubility (i.e. O2 concentration atsaturation), and the gas transfer velocity (e. g., Reuer et al., 2007).Gross O2 bioflux was calculated from Prokopenko et al. (2011,above equation).

The calculation of net and gross bioflux relies on the assump-tion that there is negligible addition of air during the collection ofthe water samples using the ship’s underway supply, e.g. fromeither air forced under the hull by ship motions or exposure to airin the pumping system. The relationship between δ18O of O2 andthe O2/Ar ratio provides an assessment of air injection duringcollection. In the presence of only biological processes, these twovariables are anti-correlated along a line that passes slightly abovethe atmospheric equilibrium point (at O2/Ar��86 per mil relativeto air, and δ18O of O2 about 0.8‰, in the Southern Ocean). Thisoccurs because the photosynthetic addition of O2 raises O2/Arwhile lowering δ18O of O2 (since photosynthetic O2 has the sameδ18O as seawater, around �24‰ with respect to air O2, which isthe standard). Conversely, respiration consumes O2, lowering O2/Ar while raising δ18O of O2 (because respiration discriminatesagainst the heavy isotope). The co-occurrence of gas exchange andvariable ratios of net/gross production complicates this correla-tion. Nonetheless, a compact relationship between δ18O of O2 andthe O2/Ar ratio was observed (as shown in Supplementary materialFig. S1). This suggests minimal addition of air to the samples, be-cause air addition changes the O2/Ar ratio while leaving δ18Onearly unchanged. Based on this diagnostic, visual inspection in-dicates that less than 5% of the samples may have experiencedsignificant air injection and only 3 points were sufficiently com-promised to require removal from the dataset (Supplementarymaterial). We also assume that there is negligible removal of O2 asseawater flows through the ship’s line from the intake to thesampling space (for a cautionary note see Juranek et al. (2010)).

Finally for this section, we note that either upwelling or netheterotrophy (Hamme et al., 2012) can cause net O2 bioflux to beless than zero. We believe that upwelling is by far the dominantcause, for two reasons. First, deep waters, undersaturated in O2,mix to the surface of the Southern Ocean, particularly south of thePolar Front. Second, in the Southern Ocean as elsewhere, there isnot enough POC in the mixed layer to sustain net heterotrophy fora large fraction of the summer half-year. We include samples withnegative bioflux in our data set (Supplementary material), but, tokeep the focus on productivity controls, we exclude biologicallyunder-saturated samples from the net bioflux values plotted in allfigures other than Fig. 3. The exception is samples that were bio-logically undersaturated by o0.3%, which is the analytical

uncertainty. In contrast, we include all gross O2 bioflux values inall figures, (regardless of their associated O2/Ar values). This is sobecause gross bioflux is much less affected by upwelling than netbioflux (since the oxygen isotopic gradient across the bottom ofthe mixed layer is relatively much smaller than the O2/Ar gra-dient). Exchange across the base of the mixed layer also impactsNCP estimates based on nutrient depletions. From model calcula-tions, the bias has been suggested to be up to 2-fold in springwhen mixed layer variability is large, and �10% in summer whenstratification is well established (Wang and Matear, 2001; Wanget al., 2001).

2.4. Additional data sources

Mixed layer macro-nutrient (nitrate, phosphate, and silicicacid) samples were collected in parallel to the oxygen samples(Fig. 2). Nutrient samples were frozen onboard and analyzed inHobart following WOCE protocols.

MLDs used in the interpretation of Astrolabe bioflux data wereestimated from XBT’s deployed by the ship. The base of the mixedlayer was identified using a temperature criterion (Tz was 0.2 °Clower than Tsurf) (Montegut et al., 2004).

PAR and Chl were characterized by remote sensing (Arrigoet al., 2008), and the extinction coefficient was parameterized interms of Chl (Morel and Maritorena, 2001). Average mixed layerPAR was calculated from the following equation:

∫= * ( − ) ( )Average ML PAR Sea surface PAR exp kz dz/MLD 4

k is the light attenuation coefficient and z is depth below thesea surface. The integral is evaluated from the sea surface to theMLD.

We also computed chl-normalized net and gross O2 bioflux, di-viding bioflux by Chl as inferred from remote sensing. Specifically, weused NASA MODIS/Aqua 2009b daily 4 km level 3 fields for remotelysensed SST, PAR, and Chl (http://oceancolor.gsfc.nasa.gov). For thepurpose of calculating chl-normalized bioflux, the in-situ biofluxobservations are matched to the nearest available satellite datapoints, and Chl is registered for the closest observation within a 1°radius from the day of collection to 3 days before and after. The in-clusion of such potentially distant matches is appropriate as thewaters we sample were likely mixed over a large area during theperiod prior to sampling. We performed sensitivity tests that suggestthat the typical maximum error in Chl due to matching of distantlocations is less than 750%, which is within the acceptable range forour application. However, the satellite chlorophyll retrievals addscatter to the calculated values of chl-normalized production. Wealso note that a number of papers discuss evidence that SouthernOcean chlorophyll retrievals are aliased toward values lower thanin situ measurements (Johnson et al., 2013; Marrari et al., 2006;Mitchell and Kahru, 2009; Szeto et al., 2011). This effect may alias ourcalculations and can be addressed when the Southern Ocean chl al-gorithm is better constrained.

The complete oxygen-based dataset is tabulated in the Sup-plementary material.

3. Results

We first present contextual observations on the spatial andseasonal variations of water column conditions. These come fromthe in-situ discrete nutrient analyses (Section 3.1) and the ARGOand satellite remote sensing estimates of mixed layer depth, in-solation, and chlorophyll (Section 3.2). We then discuss net andgross bioflux estimates (Section 3.3), and examine the relationshipbetween chl-normalized production and mixed layer irradiance.

Fig. 2. [SiO2] and [ −NO3] vs. latitude south of Tasmania during the summer half-years of 2006–2007 (top) and 2007–2008 (bottom). Red is October/November, green isJanuary, and blue is February/March. Also plotted is [ −NO3] vs. salinity.

M.L. Bender et al. / Deep-Sea Research I 112 (2016) 68–7872

3.1. Macro-nutrient drawdown

[ −NO3] and [SiO2] are plotted vs. latitude and salinity in Fig. 2 for2006–2007 and 2007–2008. The data show the well-known first-order features of Southern Ocean nutrient distributions: increasingconcentrations towards higher latitudes (as a result of upwelling),and decreasing concentrations as the growing season progresses(driven by NCP). They also reveal significant interannual variationsas well as meridional differences in the stoichiometry of nutrientdrawdown. Nitrate concentrations were near zero in Subtropicalwaters at the very northern end of the transects, and increasedstrongly across the SAZ to values greater than 20–25 mM at theSubantarctic Front. [NO�

3 ] continued to rise slowly across the PFZand AZ to reach values of nearly 30 mM south of the ACC. In con-trast, silicate concentrations were lower over a much larger por-tion of the Southern Ocean. [SiO2] was o5 mM in both Subtropicaland Subantarctic waters. Going southward, we observed bothprogressive and stepwise SiO2 increases.

Fig. 2 shows that the extent and timing of nutrient depletion, aswell as Si/N drawdown ratios, varied strongly from zone to zone.In the AZ, where winter Si concentrations were very high, the Si/Ndrawdown ratio over the season was nearly 10. In the PFZ, it wascloser to 4, and decreased further to �1 in the SAZ and to �0.1-0.2 in Subtropical waters. High ratios for the Si/N decrease in theupper water column, such as we find, were previously observedalong the nearby WOCE/CLIVAR SR3 north-south hydrographicsection (Trull et al., 2001a), and in the AESOPS study south of NewZealand (e. g., Morrison et al., 2001; Smith Jr et al., 2000). Thereare multiple contributing factors to these variations. In the AZ andPFZ, production is dominated by diatoms. Their Si uptake is knownto increase with Si availability, and the Si/N uptake ratio is ex-pected to be higher under the Fe limited conditions that prevailthere (Bowie et al., 2009; Cassar et al., 2011; Sosik and Olson,2002). The uptake ratio under Fe limitation is near 4 vs. near 1 for

Fe replete growth, at least for some species (Brzezinski et al., 2002;Franck et al., 2000; Takeda, 1998). In the SAZ and especially Sub-tropical waters, non-siliceous organisms contribute significantly toNCP, lowering the Si/N drawdown ratio (Kopczynska et al., 2001;Trull et al., 2001c). In addition to these uptake-based controls onSi/N drawdown, remineralization is important. N is recycled morerapidly than Si, attenuating the ratio of NO3

-/SiO2 drawdown(Green and Sambrotto, 2006; Rubin, 2003). N/P drawdown stoi-chiometry also revealed zonal variations, with the median uptakeratio of N/P ranging from 13 to 19 in the Subtropical, Subantarctic,and Antarctic zones, but only �9 in the PFZ. This result agrees wellwith (Lourey and Trull, 2001), who observed a ratio of 15.1 in theSAZ and 8.3 in the PFZ, but the cause remains unclear.

Calculation of NCP from the seasonal surface nitrate drawdownis difficult, because vertical distributions were not measured yetappear to be important, especially in the PFZ (Lourey and Trull,2001; Parslow et al., 2001). As well, both vertical mixing andnorthward Ekman transport provide significant nutrient resupply(DiFiore et al., 2006a; Green and Sambrotto, 2006; Rubin, 2003).Model estimates suggest that these two transport modes doublefluxes calculated from springtime nutrient drawdown (Wang andMatear, 2001; Wang et al., 2001). Keeping these caveats in mind,we make approximate first estimates as follows. The seasonal[ −NO3] decrease in the Subantarctic Zone was 3–4 mmol kg�1, andabout 2–3 mmol kg�1 between the SAF and the SACCF (Fig. 2). For asummer MLD of �40 m in the Subantarctic Zone and �80 m tothe south, and assuming C/N¼7, then seasonal NCP would beabout 1 mol C m�2 across the entire transect from the SubtropicalFront to the SACCF. In comparison, depth resolved nitrate esti-mates suggest �3–4 mol C m�2 for the SAZ from October toMarch, a period over which mixed layer depth shoaled from 200 to�50 m (DiFiore et al., 2006; Lourey and Trull, 2001).

There is considerable interannual variability in nutrient draw-down. As an example of interannual variability, October −NO3 and

Fig. 3. Hovmoller diagrams showing net O2 bioflux (sea-air biological O2 flux, mmol m�2 day�1), MLD, surface PAR, chlorophyll concentration on a log scale (mg/L), euphoticdepth, log(chl), and depth-averaged PAR vs. time for a longitude band of �30° south of Australia. Chl concentrations are from remote sensing, and MLD is from Argo floats.

M.L. Bender et al. / Deep-Sea Research I 112 (2016) 68–78 73

SiO2 concentrations were meridionally constant across the AZfrom the Polar Front (�56 °S) south to about 61 °S, in each year ofour observations. However, [ −NO3] seasonal drawdown was about1 mM in 2006–7 but closer to 3 mM in 2007–8. In this same region,Si was almost fully depleted by January in 2007 but only about halfdepleted in 2008. The timing of SiO2 depletion in our data set issimilar to other observations, which also suggest that the seasonalconsumption mainly occurs before the end of January (Green andSambrotto, 2006; Morrison et al., 2001; Rubin, 2003).

3.2. Mixed layer depth, irradiance, and chlorophyll variations

In Hovmøller diagrams, Fig. 3 shows weekly estimates formixed layer depth (MLD), euphotic depth (Z), Chlorophyll con-centrations (Chl), and surface photosynthetically available radia-tion (PAR). Also shown is depth averaged PAR for the mixed layer,and the differences between the mixed layer and euphotic depths.Also shown are our sparser measurements of net O2 bioflux, i.e.estimates of NCP.

There are clear seasonal cycles and considerable interannualvariability in all these properties. MLD follows the well-knownseasonal cycle: deep in winter and shallow in summer, The sea-sonality is at its largest in the Subantarctic Zone (Rintoul and Trull,2001; Rintoul et al., 1997; Sokolov, 2008; Trull et al., 2001a). Togeneralize from 3 years’ data, mixed layers at the beginning ofOctober are 4150 m deep north of the Polar Front, and shallowerto the south. Mixed layers shoal in October in the north and No-vember in the south. In the fall there is a slow transition back todeep wintertime values. Sea surface PAR values track the sun, withmaximum values close to summer solstice. Average values ofmixed layer PAR have greater variation in seasonal timing. In

2007–2008, mixed layer average PAR in the northern part of thestudy region was highest before summer solstice. In the next twoyears it was highest after. This pattern tracks MLD, and is a resultof interannual variability in water column stratification. ShallowerMLD’s were more prevalent in the early part of the growing seasonin 2007–2008, and more prevalent in the latter part the following2 years.

In the northern part of the study area, Chl levels rose later inthe growing season than the shoaling of MLD or the increase inmixed layer PAR. In other words, biomass (or at least chlorophyll)continued to increase after stratification. Highest Chl levels werenot observed until the end of January, or February. During summer,Chlorophyll concentrations reached a maximum around the Sub-tropical Front, then decreased towards the south. There was asuggestion that Chl levels also increased at the southern end of thetransect, as Antarctica was approached-a feature that is seen morestrongly to the east of the Astrolabe transit in satellite data (So-kolov, 2008). Even the higher Chl accumulations that occurred inthe SAZ were not large, rarely reaching more than 1 mg L�1 (3).These levels are similar to those reached in the iron-limited sub-arctic North Pacific (Harrison et al., 2004) and can be contrastedwith levels 3–5 times higher that develop in the more iron-rich(Ryan-Keogh et al., 2013) sub-arctic North Atlantic spring bloom(e.g. Mahadevan et al., 2012).

During the mid-summer periods of highest Chl accumulations,the mixed layer was generally shallower than the euphotic zone(see citations in Section 2.1). Vertical mixing would cause ourfluxes to overestimate ML production. The effect is strongest northof the subtropical front in summer (Fig. 3), and would cause us tooverestimate summertime NCP between 43–46° south. This effectis unlikely to be important south of the Subtropical Front, where

Fig. 4. Left: Net and gross bioflux vs. latitude for all samples. Right: Chlorophyll-normalized net and gross bioflux vs. latitude for all samples. Error bars correspond tostandard errors. Upper part of the plot shows the number of samples for each degree of latitude.

M.L. Bender et al. / Deep-Sea Research I 112 (2016) 68–7874

only a small fraction of primary production takes place below themixed layer (see Sections 2.1).

3.3. Spatial and temporal distributions of Gross and Net CommunityProduction

Net O2 bioflux (as an estimate of NCP) and gross O2 bioflux (asan estimate of GPP) are plotted vs. latitude and averaged over 1°latitude bands in Fig. 4. In this and subsequent discussions ofmeridional variations, we exclude from the data set results fromthe Astrolabe cruise that is to the east of the other cruise tracks(Fig. 1). Net bioflux (averaged over all samples with biological O2

supersaturation4�0.3%) was 18717 mmol m�2 day�1 (1s,n¼213). Gross O2 bioflux was 917110. As discussed above, NCP islikely to be slightly higher than net bioflux, which is aliased low byvertical mixing. The ratio of both average and median net O2

bioflux/gross O2 bioflux was 0.20. These values are all similar tothose of other studies of large domains in the Southern Ocean(Green and Sambrotto, 2006; Morrison et al., 2001; Reuer et al.,2007; Rubin, 2003).

Gross bioflux was highest from 43 to 48 °S; in contrast, netbioflux showed no strong meridional trend (Fig. 4). This situationmay be compared with that during the autumn SAZ-SENSE voyage(Feb. 2007). During this study, 14C production, gross O2 bioflux,and net O2 bioflux were all highest just to the south of the Sub-tropical Front near 45–46 °S (Cassar et al., 2011). However, it isimportant to note that the SAZ-SENSE data did not show amonotonic southward decrease of net bioflux. Net bioflux was low

SACCF AZ PFZ SAZ STZLatitude (N )

Fig. 5. Average net and gross bioflux for the major zones of the Southern Ocean south o46 °S), SAZ, Subantarctic Zone (52–46 °S); PFZ, Polar Frontal Zone (56–52 °S); AZ, Antarcti

in the southern part of the SAZ, then rose south of the Polar Front.In our Astrolabe data, gross bioflux was highest at the northernend of the transect, but not net O2 bioflux. On the other hand, datafrom the New Zealand sector (Reuer et al., 2007) showed a per-sistent southward decrease in both net and gross bioflux fromNovember or December through February.

Monthly average values (and standard errors) for gross and netbioflux are plotted vs. Southern Ocean Zone, from STZ to SACC, inFig. 5. These data again illustrate the main temporal and mer-idional trends. There is a southward decrease in gross bioflux butlittle or no decrease in net bioflux. Maximum values of net biofluxcome after the New Year. South of the Subtropical Front, maximumvalues of both net and gross bioflux occur in January. This featuremay be linked to the rapid drawdown of SiO2 in the December-January period.

The maintenance of high gross production into the beginning ofMarch may be associated with elevated chlorophyll and shallowmixed layer depths during this time (Fig. 3). Iron is likely de-creasing throughout the growing season (Bowie et al., 2009;Sedwick et al., 2008), but light and chlorophyll apparently win inour domain.

For the period from October to the beginning of March (�130days), net O2 bioflux averaged 17, 18, and 17 mmol m�2 day�1 forthe Subtropical Zone, Subantarctic Zone, and Polar Front zone re-spectively. For a growing season of 130 days and ΔCO2/ΔO2¼117/170 (Sarmiento et al., 1998), NCP inferred from our O2/Ar datawould be 1.5 mol C m�2 year�1. Thus our estimates are in rea-sonably good agreement with estimates from a detailed seasonal

SACCF AZ PFZ SAZ STZLatitude (N )

f Australia as a function of the month of sampling. STZ, Subtropical Zone (north ofc Zone (63–56 °S), and SACC, South of the Antarctic Circumpolar Current (S of 63 °S).

M.L. Bender et al. / Deep-Sea Research I 112 (2016) 68–78 75

cycle study at 46° 56′S, 142° 15′E using hourly autonomous mooredobservations of O2 and pCO2 (�2 mol C m�2) (Shadwick et al.,2015a; Weeding and Trull, 2014). Estimates from −NO3 drawdownare similar (Section 3.2). As with the Chl abundances, our NCPestimates are relatively low overall (o20 mmol m�2 day�1).These values are similar to sub-Arctic North Pacific waters(e.g. Harrison et al., 2004). They are about an order of magnitudebelow NCP values during the spring bloom in the sub-Arctic NorthAtlantic (Alkire et al., 2012; Bender et al., 1992; Quay et al., 2012).There, NCP can be as high as 1 mol C m�2 over a period of order1 week, equivalent to net O2 bioflux values of about200 mmol m�2 day�1.

The seasonal cycle of NCP inferred here for the SubantarcticZone is very different from that of Weeding and Trull (2014), whomeasured [O2] and gas tension using a mooring. They diagnosedvery high values of O2 NCP in early spring. These high values wererequired to explain the observed time series of [O2], given evi-dence for rapid entrainment from deep mixed layers and highlyvariable mixed layer depths. This entrainment kept O2 under-saturated at the surface until the middle of December, after whichtime vertical mixing has little influence on the O2 balance of themixed layer (their Fig. 5C). In contrast, the mixed layer was gen-erally supersaturated in biological O2 during our cruises (TableS-1). Also, we calculated NCP only for waters that were biologicallysupersaturated in O2. Thus biases from entrainment in our ob-servations are likely to be small.

The difference between the two studies may also reflect spatialvariations in the depth and intensity of late winter and earlyspring deep convection south of Australia. The Southern OceanTime Series site southwest of Tasmania studied by Weeding andTrull (2014) exhibits deeper winter mixed layers than along ourAstrolabe transects and the nearby WOCE/Clivar SR3 line (Donget al., 2008; Rintoul and Bullister, 1999; Rintoul and Trull, 2001).Interannual variability may also contribute to the difference. Weobserved biological supersaturation in the spring of 2007 and2008, but mixed saturation status in spring, 2009. Significant in-terannual variability in biological O2 supersaturation and net O2

bioflux is hardly surprising given the large difference in Sub-antarctic Zone nitrate concentrations in October and November2006, relative to 2007. Spatio-temporal variations may also explainhigh NCP inferred from nutrient and carbon dioxide observationsin Drake Passage in springtime (Munro et al., 2015). Overall, theseother observations of high springtime NCP do not invalidate ourconclusion that, in O2-supersaturated waters south of Australia,springtime productivity is not particularly high.

4. Discussion

4.1. Relationships between Chl, mixed layer PAR, and gross and netbioflux

The relationship between sea surface PAR, MLD, and produc-tion has been at the heart of attempts to understand seasonal andspatial variability in productivity back to at least the classic paperof Sverdrup (1953). It is also at the center of algorithms to calculate14C production from remotely observed properties of the sea sur-face (e. g., Behrenfeld and Falkowski, 1997). Here we analyze therelationship between these properties using a simple approachthat has basic mechanistic content. Our objectives are to identify asimple and transparent relationship between these properties, andto advance the possibility of simulating NCP from remotely sensedproperties and Argo/BioARGO float data.

Specifically, we focus on the relationship between average vo-lumetric production in the mixed layer normalized to Chl, andaverage irradiance in the mixed layer. One expects these terms to

be related, because productivity increases as irradiance rises andas Chl rises. Our work follows, for example, the Vertically-Gen-eralized Productivity Model of (Behrenfeld and Falkowski, 1997).Upon rearranging Eq. (10) of their paper, one gets an equationstating that the rate of Chl-normalized volumetric primary pro-duction in the euphotic zone scales with irradiance. Their irra-diance term takes into account light saturation, and their model ismore process- rich than ours (therefore also more complex), butthe basic premises are similar.

In our model, the relevant terms are calculated as follows:

( )− =

* ( )

OChl normalized volumetric gross bioflux in the ML gross

bioflux/ MLD Chl 5

2

( )− =

* ( )

Chl normalized volumetric net bioflux in the ML net O

bioflux/ MLD Chl 6

2

The equation for calculating average ML irradiance was pre-sented earlier. Chl is in mg m�3. The units of volumetric biofluxare mmol O2 mg Chl�1 day�1. Average ML irradiance has the unitsof moles photons m�2 day�1.

In evaluating the relationship between normalized bioflux andaverage ML irradiance, we exclude all samples withChlo0.10 mg m�3, because at these low levels Chl retrieved fromspace may have large uncertainties (e. g., Johnson et al., 2013). Asoutlined above, we exclude samples from the net bioflux analysisthat were biologically undersaturated in O2 by 40.3‰. We alsoexclude two measurements of gross bioflux that were extremeoutliers. We then plot normalized gross O2 bioflux and normalizednet O2 bioflux vs. average ML irradiance (Fig. 6).

Chl-normalized volumetric gross and net O2 bioflux both risewith increasing values of average ML irradiance, albeit with con-siderable scatter. Out to a light flux of �25 mol m�2 day�1, there isno strong indication that irradiance is saturating. The R2 is 0.44 forboth normalized net bioflux and normalized gross bioflux plotted vs.average ML irradiance. This level of shared variance is remarkablyhigh given errors in derived properties (satellite Chl observations, gastransfer velocity, mixed layer PAR, light attenuation coefficient) andthe simplistic nature of the assumptions for calculating O2 bioflux(steady state, no water flux across the base of the mixed layer).

We analyzed the apparent absence of light saturation in theplots of normalized production vs. average ML irradiance using asimple model based on curves of 1-h 14C production vs. irradiance(Westwood et al., 2011), measured on the SAZ-SENSE cruise. Ac-cording to Westwood et al., saturating irradiance during the SAZ-SENSE cruise averaged 56 mmol photons m�2 s�1 when evaluatedfor those stations where the MLD was within the euphotic zone.For a 16-hour day, this number corresponds to about 3 mol pho-tons m�2 day�1. It is approximately equal to the median value ofaverage daily ML irradiance for our samples (Fig. 6). However,Fig. 6 shows that gross and net O2 bioflux continue to increase athigher irradiances, despite the fact that average daily ML irra-diance is saturating. There are 2 reasons. First, light decreasesexponentially with depth. Therefore, most of the mixed layer liesbelow the depth of average ML irradiance. Second, irradiance willbe low, and production will be light-limited, towards the begin-ning and end of the photoperiod. For these 2 reasons, productionduring the photoperiod may be light-limited even if average ir-radiance in the mixed layer exceeds the light-saturated value.

In Fig. 7, we quantify these effects by plotting relative O2 pro-duction vs. irradiance for the case where saturatingirradiance¼56 mmol photons m�2 s�1. We assume that in-stantaneous production increases linearly with irradiance until itreaches the saturation value (at 56 mmol photons m�2 s�1), and isconstant thereafter. In our calculation, insolation decreases

Fig. 6. Net (left) and gross (right) O2 volumetric bioflux normalized to chlorophyll, plotted vs. average mixed layer irradiance. Net and gross O2 bioflux are calculated frommixed layer data and approximate O2 NCP and O2 GPP in the mixed layer. Normalized net bioflux is plotted only for samples where the biological O2 supersaturation is4�0.3% (within uncertainty of zero).

Fig. 7. Relative production vs. average daily irradiance in the mixed layer. Pro-duction is assumed to increase linearly up to saturating irradiance (taken as 56mmol photons m�2 s�1), and be constant at higher irradiances. This plot shows thatdaily production increases linearly with average mixed layer irradiance up to ir-radiance values much higher than the instantaneous saturating irradiance,3.2 mol m�2 day�1. The reason is that irradiance early in the day, late in the day,and lower in the mixed layer can be undersaturated even when average daily MLirradiance is much greater than the instantaneous saturating value.

M.L. Bender et al. / Deep-Sea Research I 112 (2016) 68–7876

symmetrically about the noontime value along a hyperbolic curvethat goes to zero at 4 AM and 8 PM (i.e., 16 h/8 h light/dark cycle).We calculate relative ML production for MLD from 1 to 100 m, in1 m intervals. Corresponding values of average ML irradiance de-crease from 45 to 6 mmol m�2 s�1. Mixed layer O2 productionsaturates at an average irradiance of �15 mol photons m�2 day�1,approaching the highest value observed for our samples. Hencemost of the mixed layer is light-limited for most of the day, and weexpect that gross O2 production will increase with increasing ir-radiance, given the ambient range of ML irradiances. In practice, ofcourse, production must eventually saturate at high irradiance.This dependence may be present in the curve of normalized grossO2 evolution (Fig. 7, left), but there is a lot of scatter.

Various factors complicate the relationship between light, Chl,and O2 production. Perhaps the most important is the iron

sufficiency of the ecosystem, which also influences μ. There areample data suggesting that iron sufficiency changes in the studyregion (e. g., Cassar et al., 2011). Hiscock et al. (2008) argued thatthe effect of iron on the physiology of individual phytoplankton isto increase the initial slope of the P-I curve while leaving light-saturated production unchanged. In their experiments, the slopevaried by about a factor of 2 between Fe-limited and Fe-sufficientphytoplankton. This variability would lead to scatter in the curveof gross O2 production (or bioflux) vs. average ML irradiance, butthe basic relationship would be unchanged. In fact the greatestrelevant effect of Fe is simply to cause a rise in the Chl. We im-plicitly acknowledge this effect by normalizing to Chl. Other ob-vious confounding factors include water temperature, light accli-mation, and the changing nature of Southern Ocean ecosystems.

Despite these complications, our results show that there is astrong relationship between volumetric net and gross biofluxnormalized to Chl, and average ML irradiance. These relationshipscan be exploited, using remotely sensed Chl values and MLD’s fromARGO floats and other sources, to scale local estimates of GPP andNCP to broader areas of the Southern Ocean. Scaled values can beused to achieve a better understanding of process-level controlson production, and to improve the simulation of biogeochemicalcycles in models.

4.2. Net/gross bioflux variations

There are many possible controls on the overall trend of lowerNCP/GPP ratios in the northern regions of the Southern Oceanobserved here (Fig. 4). For example, zooplankton grazing and re-mineralization are both more effective in warmer waters (Lawset al., 2000; Steinberg et al., 2008). In addition, lower levels of SiO2

in the north may limit the growth of diatom micro-plankton,whose size allows them to more readily escape grazing pressure(Assmy et al., 2013; Morel et al., 1991). Importantly, the complexityof tropho-dynamics means that high NCP can still occur in smallphytoplankton communities, as has been observed previously inthis region (Cassar et al., 2015), including contributions to carbonexport to the ocean interior (Richardson and Jackson, 2007).

5. Conclusions

Our data allow us to answer our three key research questions,outlined in the Introduction, as follows:

M.L. Bender et al. / Deep-Sea Research I 112 (2016) 68–78 77

1. In the Australian sector of the Southern Ocean there is no evi-dence of a spring bloom in the sense of very high rates ofphotosynthesis and carbon export.

2. The oxygen based GPP and NCP estimates (and the SiO2 and −NO3drawdowns) show considerable latitudinal and year to yearvariability, but with an overall trend of higher GPP in the northand little latitudinal variation in NCP.

3. Chl-normalized values of net and gross O2 bioflux scale withaverage mixed layer irradiance, albeit with a high level of scatter.This relationship provides a useful check on models of productivitybased on remote sensing, that can be evaluated further in futurestudies. The relationship also provides a basis for scaling GPP andNCP in the Southern Ocean based on Chl observed from satellitesand mixed layer depths from ARGO floats and other sources.

Acknowledgements

Sampling on Astrolabe was supported by a French-Australian re-search collaboration and involved many dedicated field personneland help from ships' officers and crew. The Institute Polaire Francaisare thanked for allowing access to the ship and for help with fieldoperations, as is the assistance provided by Drs. C. Goyet and N.Metzl. The Australian Climate Change Research Program, AntarcticClimate and Ecosystems Co-operative Research Centre, and the In-tegrated Marine Observing System funded the collection of watersamples and nutrient analyses, and the operation of underway sen-sors reported in this study. Bender, Cassar, and Jonsson gratefullyacknowledge financial support from NASA and the Office of PolarPrograms, NSF. Alan Poole and Matt Sherlock from CSIRO were fre-quently called on to ensure equipment functioned for the voyages.

Appendix A. Supporting information

Supplementary data associated with this article can be found inthe online version at http://dx.doi.org/10.1016/j.dsr.2016.02.018.

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