234Th-derived surface export fluxes of POC from the Northern Barents Sea and the Eurasian sector of the Central Arctic Ocean

Deep-Sea Research I 68 (2012) 1–11

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234

Th-derived surface export fluxes of POC from the Northern Barents Sea and the Eurasian sector of the Central Arctic Ocean ¨ rjan Gustafsson a,n, Per S. Andersson b O a b

Stockholm University, Dept. of Applied Environmental Science (ITM) and the Bert Bolin Centre for Climate Research, 106 91 Stockholm, Sweden Swedish Museum of Natural History, Laboratory for Isotope Geology (LIG), Box 50007, 104 05 Stockholm, Sweden

a r t i c l e i n f o

abstract

Article history: Received 20 December 2011 Received in revised form 23 May 2012 Accepted 28 May 2012 Available online 9 June 2012

Settling-based surface ocean export of particulate organic carbon (POC) in the western Eurasian sector of the Arctic Ocean was investigated from the marginal ice zone (MIZ) of the northern Barents Sea to the North Pole area. Upper ocean profiles of POC were combined with corresponding dissolved and particulate 234Th activities measured with a low-volume at-sea direct beta counting protocol to constrain the 234Th-derived POC export in July and August of 2001 to 6  32 mmol m  2 d  1 for the Barents Sea MIZ dropping to 2–6 mmol m  2 d  1 for multi-year-ice (MYI) covered central Arctic stations in Nansen, Amundsen and Makarov basins. Secular equilibrium between 234Th and 238U activities in intermediate to deep waters in the Amundsen Basin (n ¼ 10) demonstrated that the at-sea measurement protocol was functioning satisfactorily. There was no distinction in POC export efficiency between the MIZ and the MYI-covered interior basins with an average ratio between 234Th-derived POC export and primary production (so-called ThE ratio) of 44%. A projected increase in primary production with retreat in areal extent of sea ice is thus likely to yield increased POC sequestration in the Arctic Ocean interior. & 2012 Elsevier Ltd. All rights reserved.

Keywords: Biogeochemistry Carbon Marginal ice zone Particle settling Oden

1. Introduction Export to subsurface strata by gravitational settling of particulate organic carbon (POC) is a key mechanism in the oceanic uptake of atmospheric carbon dioxide and is a precondition for permanent sediment burial of organic carbon. This surface ocean export is subject to significant spatial and temporal variability in both quantity and composition. Regional, seasonal and interannual variation have been related to differences in primary production, extent of carbon recycling within the surface layer, and influence of allochtonous carbon input. On a global ocean scale, there appears to be a decoupling between primary production and vertical export but with definable regional characteristics (e.g., Buesseler, 1998; Dunne et al., 2007). High latitude systems, such as the Arctic Ocean, may progress rapidly through their brief vegetative season and are thus generally believed to exhibit inefficient retention and therefore export a larger fraction of the fixed carbon by vertical settling than do lower-latitude systems (e.g., Buesseler, 1998; Wassmann et al., 2004). Continental shelf regimes, a characteristic feature of the Arctic Ocean, are generally believed to also be high in both primary and export production, albeit with differences between the many Arctic shelf seas. Given

n

Corresponding author. Tel.: þ46 8 6747317. ¨ . Gustafsson). E-mail address: [email protected] (O

0967-0637/$ - see front matter & 2012 Elsevier Ltd. All rights reserved. http://dx.doi.org/10.1016/j.dsr.2012.05.014

projections of relatively large changes to the Arctic hydrography, plankton ecology and carbon system due to climate change, there is a call for greater spatial and temporal coverage in observations of upper ocean carbon export in large areas of the Arctic Ocean (e.g., Wassmann et al., 2004; Cai et al., 2010; Wassmann, 2011). Studies to date have mainly been confined to either the Barents Sea region of the extensive Eurasian shelf (e.g., Andreassen and Wassmann, 1998; Wassmann, 2002; Olli et al., 2002; Coppola et al., 2002; Wassmann, 2004; Lalande et al., 2008; Lalande et al., 2009; Reigstad et al., 2011), or to the more narrow western Arctic (North American) shelves and neighboring Canada Basin (e.g., Hargrave et al., 1994; Moran et al., 1997, 2005; Moran and Smith, 2000; Chen et al., 2003; Baskaran et al., 2003; Trimble ¨ and Baskaran, 2005; Lalande et al., 2007; Fahl and Nothig, 2007; Lepore and Moran, 2007; Cai et al., 2010). The broad picture of Arctic Ocean surface-ocean POC export that is emerging is that the studied shelf/slope waters exhibit higher export fluxes than the perennially ice-covered central basins, consistent with elevated primary production and inefficient recycling in these Arctic shelf/ Marginal Ice Zone (MIZ) regions. It is apparent that there is a shortage of information from the vast East Siberian Arctic Shelf (ESAS; Laptev and East Siberian Sea) with only a handful of ¨ observations from the outer Laptev shelf/slope (Fahl and Nothig, 2007; Lalande et al., 2009; Cai et al., 2010). For the deep basins of the interior Arctic Ocean, there exists only one study with an extensive number of observations (Cai et al., 2010). This study,

2

¨ . Gustafsson, P.S. Andersson / Deep-Sea Research I 68 (2012) 1–11 O

performed in the unusual summer of 2007 (lowest ice coverage on record), offers one of the largest and most detailed 234Th-POC data sets ever published with outer shelf–basin transects from Barents, Kara and Laptev Sea as well as a transect crossing the eastern Amundsen Basin into the Canadian Basin. Beyond this one, only a handful of 234Th-derived upper ocean POC export fluxes exist for the Central Arctic Basins (Moran et al., 1997; Baskaran et al., 2003; Trimble and Baskaran, 2005). These pioneering studies report 234 Th deficits in the deep Canada Basin and concluded that a small POC export from local biogenic production and scavenging was discernable. Given a predicted scenario due to climate warming of a large component of the interior Arctic becoming ‘‘MIZ-like’’ and the putative differences in POC export dynamics of a perennially icecovered regime and a MIZ regime (e.g., Reigstad et al., 2002; Lalande et al., 2009; Slagstad et al., 2011), it is the overarching objective of the present study to contribute further observationbased information of upper ocean carbon export in the Eurasian sector of the Arctic Ocean, spanning from the Barents Sea MIZ to the North Pole area. We here provide detailed upper ocean profiles of 238U-234Th that is combined with concurrent POC distribution and primary production from several locations in the northern Barents Sea as well as from single locations each in the deep Nansen, Amundsen, and Makarov Basins, respectively. The study thus affords coherent comparison of upper ocean POC export fluxes and export efficiencies in a key MIZ region and in interior ice-covered regimes. Results are interpreted with additional characterization of the organic matter composition and ancillary system descriptors as well as within the context of previous investigations.

2. Methods 2.1. The Swedarctic 2001 expedition Samples were collected during the Arctic Ocean 2001 (AO-01) expedition onboard the Swedish icebreaker ODEN. The AO-01 was ongoing 29 June – 28 August 2001 with areal foci of northern Barents Sea east and north of Svalbard as well as in the Eurasian Basin of the high Arctic Ocean. The expedition carried several different programs including atmospheric, biogeochemical, seismic, remote sensing and physical oceanographic components. 2.2. Sampling and at-sea direct beta counting of

234

Th

Water column samples were collected for 234Th from six standard depths in the top 100 m at each of the eight stations (Fig. 1). Twenty liter seawater samples were drawn from a 24bottle Niskin rosette (12 l sizes) into low-density polyethylene cubitainers (Cole-Parmer) held in HDPE crates. The procedure for sample treatment and direct at-sea beta counting outlined by Rutgers van der Loeff and Moore, 1999 and also described in Gustafsson et al., 2006 was employed. Briefly, the samples were immediately transported to the shipboard laboratory and vacuum filtered through 142 mm diameter polycarbonate track-etch membrane filters (1.0 mm; Nuclepore Inc.) for particulate 234Th. Ten liter of the dissolved fraction was transferred to an acidcleaned 10 l cubitainer and the 234Th was co-precipitated and scavenged with a formed MnO2 floc. The precipitate was allowed to mature for 24 h after nucleation and then filtered through the same type of membrane filters as described above. In a HEPAfiltered laminar flow bench, all filters were dried and then folded six times into a 12  12 mm 64-layer thick square, wrapped in one layer of thin polyethylene foil (to filter out potentially interfering weaker beta emitters such as 40K; see also Rutgers

180°

12

11 85°N

6 5 1

80°N

4

40°E 3 0°

2

20°E

Fig. 1. Bathymetric map of the northern Barents Sea and the Eurasian sector of the central Arctic Ocean illustrating the locations of the studied sampling stations of the Swedarctic Arctic Ocean 2001 (AO-01) expedition.

van der Loeff and Moore, 1999) and assembled on speciallyconstructed plastic stubs. The stubs were then inserted into a low-background five-channel beta counter (Riso¨ GM 25-5) for direct at-sea counting onboard the ODEN. The inherent detector backgrounds at-sea (with empty stubs) were 0.18–0.21 countsper-minute (cpm) for the five different channels. Each of the 300 filters counted as part of the AO-01 campaign was recounted after4 6 half-lives of 234Th (t1/2 ¼24.1 day) to yield an average total matrix background of 0.29 cpm. The counting efficiencies were calibrated for these open-sea samples by preparing FeOH3 precipitates of a U standard. Five such standards were prepared according to the protocol outlined by Rutgers van der Loeff and Moore, 1999 on the same filter types used for the samples. The self absorption in this system is due to the filter material. The counting efficiencies of the five beta detector channels were frequently constrained using this set of five 238U-234Th standards. It was found for our equipment to be in range of 0.373–0.383. Detector-channel specific backgrounds and efficiencies were applied to the individual sample counts, individual sample matrix backgrounds were subtracted and the resulting 234Th activities were corrected for U ingrowth and decay to the time point of sample collection. The 238U activities were estimated from the linear relationship with salinity from Chen et al., (1986). Salinities were obtained with a Sea Bird 911 CTD instrument mounted on the 24-bottle rosette and calibrated during the expedition with a ¨ et al. 2002). The 234Th Guildline 8400B Autosal salinometer (Bjork activities resulting from applying this method to mid- and deep waters in the Amundsen Basin were compared with expectations from secular equilibrium with 238U to independently test for the quantitative recovery of this direct beta counting procedure (see Results below). 2.3. Sampling and analysis of particulate organic matter Samples of 2–4 l for POC were drawn at the same time as for Th from parallel Niskin bottles. The polycarbonate sample bottles were immediately vacuum filtered in the shipboard laboratory using pre-combusted 47 mm borosilicate filters 234

¨ . Gustafsson, P.S. Andersson / Deep-Sea Research I 68 (2012) 1–11 O

(  0.7 mm; GF/F; Whatman Inc.) held in a stainless steel filter holder. Maximum filter pressure was about 0.5 bar to prevent significant cell lyzing. The filters were stored at  18 1C until further treatment in the Stockholm University laboratory. The GF/F filters were then sub-sampled by punching out 4 mm diameter circles, which were transferred to pre-combusted Ag capsules and subjected to in situ 1 M HCl microacidification to remove any carbonates (e.g., Gustafsson et al., 1997) prior to determination of organic carbon amount and its stable carbon isotope composition with a Europa Scientific isotope ratio monitoring mass spectrometer (UC Davis Stable Isotope Facility). Subsamples for determination of POC/234Thpart also on particles collected in sediment traps, deployed at the same location, subsamples were kindly provided for several stations by Kalle Olli and Paul ¨ Norway). Wassmann (Norwegian School of Fisheries Science, Tromso, These sediment traps were surface anchored to ice floes and consisted of parallel cylinders (7.2 cm diameter; aspect ratio 6.25) mounted in a gimballed frame equipped with a vane (KC maskiner og laboratorieudstyr, Denmark) as detailed in e.g. Olli et al. (2007). For these short (24 h) deployments, poison was avoided and there were no corrections for POC dissolution into DOC. The deployment depths of the traps used in this study are indicated in Fig. 5.

3. Results and discussion

3

16–37 m, with an average of 24 m. A strong upper halocline is perennial to this system and normally reported to be located in the upper 30 m (e.g., Wassmann et al., 2004). Consistent with the ice data, the MIZ stations 2–4 east of Svalbard had progressed further in the vegetative season and were in the spring bloom phase with high primary production (300–400 mg C m  2 d  1; Sobek et al., 2006a) and POC concentrations (Table 2). The open water station north of Svalbard (station 1) was clearly in Atlantic water and also station 5 was under influence of Atlantic water carried by the West Spitsbergen Current, as indicated by hydrology, low humic substance (HS) abundance and depletion of silicates (Sobek et al., 2006b). In contrast, the three interior basin stations all exhibited lower primary production (o 200 mg C m  2 d  1 at station 6 ando100 mg C m  2 d  1 at stations 11 and 12) and high nutrient concentrations, which resulted in overall lower POC concentrations (Table 2). Fluorescence-based concentrations of HS proved to be quite useful in distinguishing the three different regimes with valueso0.1 mg QSE/L (quinine sulfate equivalents¼QSE) in Atlantic-influenced waters, 0.3–0.5 mg QSE/L at stations just north of Svalbard (including stn 6) while the two highArctic stations (and others reported by Sobek and Gustafsson, 2004) had HS in range 4–6 mg QSE/L. The high HS content observed in the high Arctic was concomitant with a freshening of the surface layer and is likely signaling the influence of Lena River and possibly other ESAS rivers as part of the Transpolar Drift (Carmack, 1986).

3.1. General description of the study systems 3.2. Methodological considerations in estimating POC export fluxes There were large variations in both hydrological and biogeochemical parameters not only between MIZ and high-Arctic regimes but also within the Barents MIZ region (Fig. 1; Table 1). A detailed account of each system has been provided in earlier publications, including ice status, ice-rafted debris and surface water masses (Gustafsson et al., 2005), phytoplankton— zooplankton speciation and biomass (Sobek et al., 2006a; Olli et al., 2007), total organic carbon (TOC) and its optical properties (Sobek et al., 2006b), stratification, nutrients and primary production (Sobek et al., 2006a,b; Olli et al., 2007). Hence, only a synopsis of the full system descriptions is provided below. For the MIZ region east of Svalbard, stations 2 and 3 were dominated by first-year ice (FYI). Station 4 was located in open water but had been ice-covered only two weeks before the arrival of the expedition. For the MIZ region on the shelf north of Svalbard, station 1 was in open water while station 5 was covered by heavy ice. The interior Arctic Ocean deep-sea stations 6 (Nansen Basin), 11 (Amundsen Basin) and 12 (Makarov Basin) were all located in dense packs of overwhelmingly multi-year ice (MYI). Inspection of CTD profiles revealed strong stratification ranging from

The two dominating techniques for estimating surface ocean carbon export are shallow sediment traps and the 234Th proxy. Both techniques have their own limitations and advantages and a recent post-JGOFS international working group of the Scientific Council for Oceanographic Research (SCOR WG 116—Assessment of techniques for estimating surface ocean export fluxes) concluded that it is at present not possible to recommend one over the other as both are providing useful and, by enlarge, complementary information (Buesseler et al., 2006, 2007). Estimates from upper ocean sediment traps are questioned based on hydrodynamic sorting against slow-settling aggregates (Gust et al., 1996; Gustafsson et al. 2004; Buesseler et al., 2007) and issues with particle dissolution and intrusion of zooplankton ‘‘swimmers’’ (Steinberg et al., 1998; Buesseler et al., 2007), while the 234Th-derived POC export estimates are challenged by uncertainties in estimates of the POC/234Th on the settling particles (e.g., Gustafsson et al., 2006; Buesseler et al., 2006; Lalande et al., 2008) and the choice of integration depth (Moran et al., 2003). The present study benefits from extensive previous Arctic work with both sediment traps and

Table 1 Physical and hydrological description of occupied stations. Station

System

Surface water mass

Date 2001-m  d

Latitude (N)

Marginal Ice Zone AO-01-1 AO-01-2 AO-01-3 AO-01-4 AO-01-5

MIZ/ MIZ/ MIZ/ MIZ/ MIZ/

Atlantic water, MIZ Arctic water Arctic water Arctic water W. Spitsbergen Current, Atlantic water

07–13 07–05 07–06 07–07 07–11

80 77 78 78 81

High Arctic AO-01-6 AO-01-11 AO-01-12

Nansen Basin Amundsen Basin Makarov, Canadian Basin

Interior Arctic water Interior Arctic water Transpolar Drift

07–09 08–16 07–31

82 02.260 88 17.42 87 54.810

N. N. Atlantic Ocean Barents Sea Proper Barents Sea Proper Barents Sea Proper N. Barents Sea

28.370 50.240 20.980 39.160 19.750

Longitude (E)

15 29 27 33 24

49.770 51.460 18.800 08.400 32.410

25 48.050  004 32.42 154 25.200

Depth (m)

Multi-year ice (%)1

462 258 273 293 129

Open water o 10 o 10 Open water 70–80

3707 4396 3917

80–90 90–100 90–100

1 This is the fraction multi-year ice (defined as having survived at least one melt season) of the total ice coverage (the remaining is first-year ice). All ice stations were in dense pack ice (490% total ice coverage).

¨ . Gustafsson, P.S. Andersson / Deep-Sea Research I 68 (2012) 1–11 O

4

Table 2 Detailed data for

238

U,

234

Th, and POC in the upper water column (0–100 m) obtained during AO-01.

Station

Depth (m)

238

U (dpm L  1)

234

Th diss. (dpm L  1)

AO-01-1

5 10 20 40 60 100

2.398 2.415 2.431 2.468 2.476 2.471

1.426 7 0.036 1.715 7 0.041 1.628 7 0.039 1.833 7 0.044 1.956 7 0.055 1.961 7 0.053

AO-01-2

5 10 20 40 60 100

2.347 2.359 2.411 2.432 2.435 2.441

AO-01-3

5 10 20 40 60 100

AO-01-4

234

POC (lM)

0.259 7 0.007 0.252 7 0.007 0.312 7 0.009 0.166 7 0.005 0.178 7 0.006 0.210 7 0.007

1.685 70.062 1.967 70.073 1.940 70.072 1.999 70.082 2.134 70.094 2.171 70.091

27.57 4.6 22.87 0.7 20.77 0.9 15.67 0.1 9.37 0.3 7.37 0.4

0.9397 0.035 1.2807 0.046 1.2087 0.046 1.269 7 0.046 1.734 7 0.055 1.897 7 0.061

0.469 7 0.013 0.498 7 0.013 0.629 7 0.019 0.166 7 0.007 0.157 7 0.006 0.155 7 0.009

1.408 70.065 1.778 70.080 1.837 70.089 1.435 70.081 1.891 70.094 2.051 70.131

23.57 0.1 20.77 2.6 21.37 1.0 3.57 0.4 3.47 0.0 3.27 0.5

2.339 2.366 2.398 2.431 2.433 2.435

1.771 7 0.058 1.414 7 0.047 1.877 7 0.056 1.763 7 0.055 0.8307 0.025 1.0767 0.040

0.327 7 0.010 0.393 7 0.012 0.515 7 0.014 0.284 7 0.008 0.286 7 0.010 0.153 7 0.005

2.097 70.096 1.807 70.081 2.392 70.097 2.047 70.087 1.11 70.052 1.229 70.063

10.67 0.2 10.87 1.6 25.67 0.1 5.37 1.3 1.97 0.2 1.57 0.1

5 10 20 40 60 100

2.315 2.315 2.425 2.434 2.437 2.442

1.5907 0.060 1.493 7 0.058 1.6607 0.061 1.872 7 0.067 1.322 7 0.048 1.169 7 0.044

0.291 7 0.011 0.293 7 0.011 0.490 7 0.016 0.372 7 0.014 0.293 7 0.012 0.134 7 0.007

1.880 70.100 1.786 70.098 2.150 70.105 2.244 70.117 1.615 70.089 1.303 70.085

18.87 0.5 16.87 3.7 14.47 1.1 7.27 0.0 5.17 0.0 1.67 1.8

AO-01-5

5 10 20 40 60 100

2.301 2.320 2.392 2.459 2.465 2.465

1.9207 0.054 1.7707 0.051 1.793 7 0.050 2.183 7 0.052 1.969 7 0.049 1.736 7 0.045

0.059 7 0.004 0.081 7 0.003 0.172 7 0.005 0.283 7 0.014 0.245 7 0.013 0.346 7 0.014

1.979 70.153 1.851 70.096 1.966 70.082 2.466 70.135 2.214 70.130 2.082 70.101

14.47 2.4 11.67 1.2 8.37 6.8 8.87 2.2 7.77 0.2 5.67 0.1

AO-01-6

5 10 20 40 60 100

2.400 2.410 2.419 2.432 2.437 2.461

1.999 7 0.076 2.355 7 0.085 2.163 7 0.047 2.167 7 0.082 1.122 7 0.054 1.433 7 0.063

0..036 7 0.002 0.034 7 0.002 0.027 7 0.003 0.039 7 0.002 0.018 7 0.001 0.1007 0.004

2.035 70.121 2.389 70.141 2.190 70.166 2.206 70.139 1.140 70.081 1.533 70.090

3.07 0.3 3.87 0.4 3.27 0.1 7.47 3.5 1.17 1.1 1.57 1.3

AO-01-11

5 10 20 40 60 100

2.274 2.274 2.278 2.374 2.391 2.412

2.142 7 0.066 2.0867 0.067 2.258 7 0.070 2.4047 0.072 2.4407 0.073 2.4407 0.049

0.044 7 0.003 0.055 7 0.004 0.019 7 0.001 0.028 7 0.002 0.027 7 0.002 0.0307 0.002

2.187 70.179 2.142 70.159 2.277 70.170 2.432 70.201 2.467 70.172 2.471 70.140

n/a 5.27 0.3 3.07 0.1 2.07 0.8 2.07 0.1 0.97 0.0

AO-01-12

5 10 20 40 60 100

n/a 2.211 2.211 2.243 2.344 2.408

n/a 1.622 7 0.065 1.689 7 0.068 1.827 7 0.080 2.0737 0.087 1.933 7 0.081

n/a 0.088 7 0.005 0.091 7 0.005 0.093 7 0.005 0.083 7 0.005 0.045 7 0.003

n/a 1.709 70.115 1.779 70.124 1.919 70.141 2.156 70.149 1.978 70.165

n/a 4.47 0.0 2.77 0.1 2.27 0.1 1.47 0.3 1.17 0.2

the 234Th approach. It extends the observational coverage of 234Thderived POC export fluxes with a July-August 2001 expedition to the northern Barents Sea MIZ and sites in the Eurasian sector of the deep Arctic Ocean. 3.3.

234

Th activities, distribution and fluxes

Dissolved, particulate and total 234Th as well as total 238U activities in the northern Barents Sea and in the interior Eurasian Arctic Ocean are listed in Table 2. The activity-depth profiles reveal at least two different features (Fig. 2). First, all northern Barents Sea stations show a 238U-234Th disequilibrium in the top 40 m surface layer with 234Th/238U activity ratios in the MIZ on the shelf east of Svalbard in the range 0.6-1.0 and in the Atlanticinfluenced MIZ-shelf waters north of Svalbard in the range 0.7-1.0. The interior Arctic Ocean reveals smaller but discernable disequilibria. The 0-40 m layer exhibit 234Th/238U activity ratios in

234

Th part. (dpm L  1)

Th total (dpm L  1)

southern Nansen Basin of 0.85-0.99, in 881N Amundsen Basin a mere 0.94-1.02 while the 881N site on the other side of the Lomonosov Ridge in the northern Makarov Basin yields 0.77–0.92. These activity ratios for Barents Sea are broadly in a range similar to what has previously been reported for both southern-central Barents Sea (Coppola et al., 2002) and northern Barents Sea (Lalande et al., 2008; Cai et al., 2010). Lalande and co-workers demonstrated that it is important that we collect information for different years as they found activity ratios of 0.8 in July 2003 but only 0.5–0.6 in May 2005 (both for top 60 m). For the North Pole area, Moran et al. (1997) reported similar activity ratios as was observed in the current study, whereas Cai et al., (2010) only found slight disequilibrium (0.9–1.0) and only for top 25 m. Taken together, the depth extent of these surface water 238U 234 - Th disequilibria are consistent both with the mixed layer depths and settling export mediated by either autochthonous biogenic particles or release of ice-rafted allochthonous particles.

¨ . Gustafsson, P.S. Andersson / Deep-Sea Research I 68 (2012) 1–11 O

Activity (dpm/L)

Activity (dpm/L) 0

0.5

1.0

1.5

5

2.0

2.5

3.0 0

0.5

1.0

1.5

2.0

2.5

3.0

3.0

20

depth (m)

40

60

80 part-Th diss-Th

100

sum-Th tot-U 0 120 0

0.5

1.0

1.5

2.0

2.5

3.0 0

0.5

1.0

1.5

2.0

2.5

0

0.5

1.0

1.5

2.0

2.5

3.0 0

0.5

1.0

1.5

2.0

2.5

0

0.5

1.0

1.5

2.0

2.5

3.0 0

0.5

1.0

1.5

0 20

depth (m)

40

60

80

100

120 3.0

0 20

depth (m)

40

60 80

100

120 2.0

2.5

3.0

0

20

depth (m)

40

60

80

100

120 Fig. 2. Upper ocean depth profiles of the activities (dpm/l) of total 238U (open red squares), particle-associated 234Th (green diamonds), dissolved 234Th (blue circles) and the sum of particulate and dissolved 234Th (black squares). The uncertainty bars indicate one standard deviation of propagated errors based on counting statistics. Station number and water mass is indicated above each panel. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

¨ . Gustafsson, P.S. Andersson / Deep-Sea Research I 68 (2012) 1–11 O

A second feature in our (and many others’) activity-depth profiles is a subsurface increase in the 238U -234Th disequilibria at several stations in the Svalbard shelf area. This feature is particularly salient in the 60–100 m depth interval at stations 3–6 and is likely to reflect scavenging of 234Th by particles resuspended from shelf sediments. The observation of this subsurface scavenging also at the deep-water station in southern Nansen Basin (stn 6) is consistent with off-shelf nepheloid transport. We believe this process is similar to the observations of subsurface scavenging due to off-shelf particle export that has been investigated in greater detail on the outer Chukchi shelf—Canada Basin (e.g., Moran et al., 1997, 2005) and Beaufort shelf—Canada Basin (Moran and Smith, 2000; Baskaran et al., 2003; Trimble and Baskaran, 2005) on the other side of the Arctic Ocean. At the stations where we observe this subsurface phenomenon the activity ratio is 0.5–0.6 over the 60– 100 m depth interval. Similarly, Moran and co-workers observe in repeated detailed transects crossing the Chukchi slope that the shelf-scavenging feature is centered around 100 m and their activity ratios are generally 0.5–0.7 (e.g., Moran et al., 2005). This off-shelf particle transport process is apparently active over the time scale of 234Th (Eup to 3 months). These workers documented the feature up to 30 km off the Chukchi shelf break, which is consistent with a maximum length scale of 100 km predicted by Moran et al. (1997) based on ambient advection and the time it would take for 234Th to regain secular equilibrium with 238U and thus erase the isotopic signal. Our Nansen Basin station 6 is within this length scale of influence from the north Svalbard shelf break (Fig. 1). Finally, our stations with subsurface maxima in disequilbrium also exhibit a larger particle fraction of the total 234Th activity at 60–100 m depths, without any apparent increase in POC (Table 2), consistent with a presence of resuspended OC-poor sediment particles (Fig. 2). It is also plausible that the subsurface increase in disequilibrium is reflecting methodological issues with insufficient 234Th yield in a lower particle regime. To test this hypothesis, and at the same time test the functioning of the at-sea direct beta counting method, we analyzed a set of samples from mid- and deep waters at the 881N Amundsen Basin station (Fig. 3). The 234Th/238U activity ratios in the ten seawater test samples were indistinguishable from unity. This result bestows credence to the integrity of the data set in this study. It is important to carefully consider over what depth the 238 U-234Th disequilibrium should be integrated to derive meaningful flux estimates. As reviewed by Moran et al. (2003), improper choice of integration depth adds to the compounded

1.2

100 m

300 m

1000 m

uncertainty of 234Th -based estimates of upper ocean export fluxes. In previous Arctic Ocean applications of the 234Th proxy, the employed integration depths have varied between 20 and 100 m, frequently without a clear basis for the choice. If the objective is to estimate particulate settling fluxes out of the surface ocean, the choice of integration depth ought to be based on either a hydrographic boundary of the surface layer (i.e., stratification; pycnocline) or a physical definition of the regime where biological particle production is occurring (i.e., the euphotic zone; the 1% light level). In this study, we scrutinized the density profiles (sigma-theta data) for each station and estimated the mixed-layer depth (mld). The mld varied from 16–37 m in the Barents shelf MIZ and over 22–30 m in the Interior Arctic stations. Since we are interested in the upper ocean particle export fluxes across a border below these depths and since any signs of influence from particles resuspended from shelf sediments only occur at depths of 60 m and below, an integration depth of 40 m was chosen for this study. For comparison, Moran et al. (2005) selected an integration depth of 50 m in their Chukchi—Canada basin study to be slightly below an estimated euphotic zone depth of 37 m. Assuming a steady-state, 1-D vertical export model, we may integrate the 234Th deficit over the 0–40 m to yield 234Th fluxes on the order of 300–900 dpm m  2 d  1 in the MIZ (Fig. 4). The 234Th fluxes in the high Arctic MYI-covered stations were lower with 240 dpm m  2 d  1 in southern Nansen Basin and 50—100 dpm m  2 d  1 in the northern reaches of Amundsen and Makarov Basins. The assumption of steady-state may result in an underestimate of the true flux when scavenging and particle export is on the increase and an overestimate when the particle export is in a decreasing phase, such as during different phases of a phytoplankton bloom (e.g., Buesseler et al., 1992; Moran and Buesseler, 1993). In this study, samples were collected from stations that most likely were in different phases of blooms and ice-melting. However, as in all previous Arctic Ocean 234Th proxy studies, the AO-01 did not provide sufficient time-series data to constrain any non-steady state changes in particle export. Transport across horizontal gradients in particle export is also assumed insignificant in the 1-D model. Previous studies on coastal environments have indicated that horizontal dispersion and advection is only significant for the 234Th mass balance under extreme situations such as high-intensity export periods and in narrow straits (e.g., Gustafsson et al., 1998; Charette et al., 2001). Next, we seek to apply the constrained 1-D 234Th export to estimate POC settling export fluxes.

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0.0 Fig. 3. The activity ratio of total 234Th to total 238U in ten samples from intermediate to deep waters at 88 N in the Amundsen Basin. The data demonstrates secular equilibrium and thus suggest full yield of 234Th in wet chemistry steps as well as functioning of the employed at-sea direct beta counting protocol.

Fig. 4. The 234Th flux in both the northern Barents Sea Marginal Ice Zone (MIZ) system as well as in the high Arctic regime. Uncertainty interval reflects one standard deviation of the propagated analytical error for both 234Th and POC.

¨ . Gustafsson, P.S. Andersson / Deep-Sea Research I 68 (2012) 1–11 O

3.4. POC distribution The depth profiles of POC concentrations demonstrated the highest values in the surface mixed layer at all stations (Table 2). For the five stations on the northern Barents Sea shelf the surface POC was in the range 12–28 mM, which is similar to surface POC reported for corresponding summer regimes on the western Arctic shelves (e.g., Baskaran et al., 2003; Bates et al., 2005; Davis and Benner, 2005). The three deep-basin stations exhibited markedly lower surface layer POC in the range 2–5 mM (Table 2), which also is similar to summertime POC reported for surface water in the deep Canada Basin (e.g., Moran et al., 1997; Wheeler et al., 1997; Baskaran et al., 2003; Bates et al., 2005). The 100 m depth POC values in the interior stations were around 1.0 mM. For all shelf and deep-basin stations, there was a marked decrease below 40 m, consistent with the stratification and putatively also the euphotic zone. These depth distributions translate into POC inventories in the top 50 m of 7–9 g m  2 for the MIZ stations located in Arctic Water and 1.8–3.9 g m  2 for the high Arctic stations. Inventories integrated down to 100 m were about 20– 30% higher on the shelf and up to 50% higher in the central Arctic, which compares well with the 0–100 m POC inventories reported by Wheeler et al. (1997) for several Central Arctic Ocean stations sampled in 1994. 3.5. POC:

234

Th ratio

Accurately constraining the POC:234Th ratio on the particles that are leaving the surface ocean is likely one of the largest uncertainties in application of the 234Th proxy method for estimating upper ocean POC export. The steady-state one-dimensional model for the 234Th-derived POC flux is: Z z ½POCz F POC,z ¼ lU ðf238 U tot gf234 T htot gÞdzU 234 ð1Þ f T hpart gz 0 where FPOC,z is the sinking flux of organic carbon (mmol m  2 d  1) crossing the export depth z (m; here selected as 40 m), l is the radioactive decay constant for 234Th (0.0288 d  1), {238Utot} and {234Thtot} is the total radioactivity concentration of 238U and 234Th (dpm m  3), integrated over the surface ocean depth at the base of which settling export is considered. [POC]z and {234Thpart}z is the particulate organic carbon concentration (mmol m  3) and particulate 234Th activity concentration (dpm m  3) of flux-weighted settling matter at depth z. Naturally, we desire to constrain the ratio of truly settling particles passing at 40 m. Techniques are now available for isolating particles based on gravitational settling, without hydrodynamic artifacts associated with traditional sediment traps, such as SPLITT (split flow-thin cell fractionation) and neutrally-buoyant sediment traps (NBST). Initial limited applications to constrain POC/234Th of the spectrum of flux-weighted settling particles using SPLITT (Gustafsson et al., 2006) and NBST (Buesseler et al., 2006) indicate that these techniques yield somewhat higher ratios than traditional sediment traps, presumably due to better ability to collect the slowest-settling particles. In the present Arctic Ocean study, a SPLITT-constrained POC/234Th ratio was obtained at the 881N Amundsen Basin station. However, as in all previous 234Th -derived POC export studies in the Arctic Ocean, the present work has to rely on approximations of the true ratio using filters and sediment traps. Since the previous, more method-centric, SPLITT study of POC/234Th indicated that filter collections of the bulk particle pool using 1 mm filters gave POC/234Th ratios closest to those from SPLITT fractionation (Gustafsson et al., 2006), the current study is based on this filtration strategy.

7

At several stations, we were also able to compare POC/234Th ratios from filters and cylindrical upper ocean sediment traps. Multiple depth profiles of POC/234Th from filters, sediment traps, and the one SPLITT fractionation are shown in Fig. 5. The filter-based POC/234Th depth profiles are geochemically consistent with overall decreasing ratios with depth signaling preferential degradation of particulate organic matter during settling. The POC/234Th for 40 m at station 6 appears anomalously high and could not be supported by ancillary data; for this station the ratio from 60 m was instead employed. There is an overall encouragingly good agreement in POC/234Th ratio between filters and sediment traps (comparison at five stations; Fig. 5) around the 40 m export horizon. This stands in contrast to the situation during the field campaigns of Lalande et al. (2008) where the POC/234Th ratios from filters were frequently factors of 2–4 higher than from floating sediment traps, presumably due to presence of the mucus from the Prymnesiophyte Phaecystis pouchetti, known to produce non-sinking POC. Fortunately, the 2001 system was dominated by diatoms and other phytoplankton (e.g., Sobek et al. 2006b; Olli et al., 2007) and thus a much better agreement in POC/234Th ratios between filters and traps (Fig. 5). Given this and the availability of filter ratio from all stations, the present study will use the filter-obtained ratios for the POC export estimations. The profiles also demonstrate that the ratios for stations from similar regimes, including the repetitions during the severalweeklong drift at station 12, tended to converge with better agreement at greater depths. Finally, the single SPLITT-derived POC/234Th ratio obtained in the surface water at the 881N Amundsen station was in excellent agreement with the ratio from filters (Fig. 5D). Hence, all necessary parameters have now been constrained for applying Eq. 1 to estimate the POC settling fluxes. 3.6. POC export fluxes POC export fluxes calculated using eq. 1 thus represent the downward POC flux through the 40 m depth horizon. The fluxes calculated for the AO-01 cruise indicate some spatial variability (Fig. 6). The northern Barents Sea MIZ stations exhibited POC export fluxes in the range 6–18 mmol m  2 d  1. The one station in the open northern north Atlantic water (station 1) was 32 mmol m  2 d  1. Lower fluxes were estimated for the Interior stations with the Nansen Basin (station 6) at 7 mmol m  2 d  1 while the high-Arctic 881N Amundsen (station 11) and 881N Makarov (station 12) were lower in POC export flux and in the range 2–3 mmol m  2 d  1 (Fig. 6). These fluxes correspond to 30– 90 day residence times of POC in the surface 40 m with respect to settling export and an average net settling (piston) velocity of about 1 m d  1. These values are in the same range as estimated from a 2-year time-series study in the central Baltic Sea (Gustafsson et al., 2004). These 234Th-constrained POC export fluxes for the northern Barents Sea MIZ and the Eurasian sector of the high Arctic Ocean may also be compared with other previous studies. A sediment trap-derived POC flux at 50 m obtained during the same AO-01 expedition for the 881N Amundsen Basin station (Olli et al., 2007) were within the uncertainty (1 standard deviation) of our 234Th derived POC export flux for that location. These POC export fluxes also compare with a summer study in the MIZ of central Barents using 24-h deployments of drifting sediment traps, reporting POC fluxes at 50 m of about 20 mmol m  2 d  1 (Wassmann and Olli, 2004), which is near the upper end of the fluxes we recorded in the northern Barents Sea MIZ. Coppola and co-workers (2002) reported 234Th-derived POC export from the Eurasian margin of the Arctic Ocean and found fluxes of 8–13 mmol m  2 d  1 for three locations in the central

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8

Fig. 5. The POC:234Th ratio on particles collected with filters, upper ocean sediment traps and, for one station (AO-01-11), SPLITT fractionation (gravitoids settling 41m/d; see e.g. Gustafsson et al., 2006). Uncertainty interval reflects one standard deviation of the propagated analytical error for both 234Th and POC.

35

Open Atl. W.

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Fig. 6. 234Th-derived POC flux in both the northern Barents Sea Marginal Ice Zone (MIZ) system as well as in the high Arctic regime. The circled data is unique as it reflects open North Atlantic water mass. Uncertainty interval reflects one standard deviation of the propagated analytical error for both 234Th and POC.

Barents Sea, which is within the range of our measurements from the northern Barents MIZ region. The reported 234Th-derived POC export fluxes on the North American Arctic shelves span a factor of 50 and are thus more difficult to compare with (Moran, 2004). More recently, Lalande et al. (2008) and Cai et al. (2010) both reported 234Th-derived POC export fluxes from Barents and also other Eurasian shelves, with widely different averages of  45 mmol m  2 d  1 and  3 mmol m  2 d  1, respectively. The Lalande et al. (2008) values may be inflated due to enhanced filter-based POC/234Th ratios due to non-sinking Phaecystis; the trap-derived POC fluxes in that same study were on average  13 mmol m  2 d  1, which is in the range of the current and other studies. For the deep interior Arctic basins, some previous studies have reported, like the current one, lower surface ocean POC export fluxes. For instance, Moran and co-workers reported 234Thderived POC export for the Canada Basin north of Chukchi Sea in the range 0.4–7 mmol m  2 d  1 while Moran and Smith (2000) and Baskaran et al. (2003) have reported fluxes for the Canada Basin off the Beaufort Sea of 2.5 and 6 mmol m  2 d  1, respectively. Finally, Moran et al., 1997 reported a single 234Th-derived

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POC export flux very close to the location of our Makarov Basin station of 1.4 mmol m  2 d  1, which compares well with the constraint from the AO-01 campaign of 2.570.2 mmol m  2 d  1. In contrast, the large set of observations in summer 2007 of Cai et al. (2010) arrive at very small POC export fluxes for the Interior Basins of on average only 0.2 mmol m  2 d  1. With the growing database of upper ocean POC export fluxes a more coherent picture is emerging with most Barents Sea MIZ fluxes in the range 10–20 mmol m  2 d  1 whereas the MYI-covered Interior Arctic Ocean seem to exhibit summertime export fluxes that are 4–10 times less and in the range 1–5 mmol m  2 d  1. It is important to note that the Cai and co-worker study documents that even lower POC export fluxes existed in late summer 2007. 3.7. Comparison of 234Th-derived POC export with primary production: ThE ratios One important objective of the present study was to evaluate the organic carbon balance in the upper ocean particularly with respect to the export efficiency in the Eurasian Arctic sector. To this end, we evaluated the relationship between POC export fluxes and rates of primary production (Fig. 7). Rates of primary production were determined during AO-01 using in situ 14C incubations for 12–24 h (e.g., Sobek et al., 2006a; Olli et al., 2007). Since our export estimates are based on the 234Th proxy, these 234Th -derived export to primary production ratios are termed ThE ratios after Buesseler (1998). Derived ThE ratios for this Northern Barents and Eurasian high Arctic study were overall high with an average of 44% (range 16-75%) of primary production being exported (Fig. 7). The interior stations (southern Nansen and 881N stations in Amundsen and Makarov Basin) had ThE ratios of 43, 33, and 27%, respectively. However, it was not possible to establish that there was a statistically resolvable difference in ThE ratios between the MIZ and high Arctic regimes or any other biogeochemical parameter. In comparison, a global assessment of ThE ratios demonstrated that there are regional/ regime differences with center of oligotrophic gyres generally showing ThE ratios of 2–10% while continental shelves on average have ThE ratios of about 25% (Buesseler, 1998). Moran and co-workers (2005) reported average summertime ThE ratios for Chukchi shelf, slope and basin stations of 32% (range 10–60%). Cai et al. (2010) reported ThE ratios of only 6% for the interior

234Th-derived POC export (mmol/m2/d)

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9

Arctic, but their estimate must be regarded as uncertain since it applied a set of primary production measurements from the 1990s to the unusual Arctic sea-ice summer of 2007. With reasoning based on the presence of overwintering zooplankton in the Arctic Ocean, Wassmann (2004) argues that ‘‘the basic fate of newly produced biogenic matter in the Arctic is retention rather than export’’, but he was also underlining the dearth of data. In a 3-week study of the AO-01 station 11 system, Olli et al. (2007) employed sediment traps and plankton ecological parameters to also suggest that the system is operating in a top-down mode with zooplankton grazing significantly attenuating POC export. Our study constrains a ThE ratio for this station of 33%. While recognizing that the coupled export efficiency data existing for the Arctic Ocean at present is limited, the available ThE ratios are suggesting that both Barents and Chukchi MIZ shelves as well as the interior MYI-covered Arctic Ocean are, in a global ocean perspective (e.g., Buesseler, 1998), rather poorly recycling and may thus be characterized as export regimes. This view is further consistent with several other high-latitude regimes also exhibiting high ThE ratios (Buesseler, 1998). Presumably the high export efficiency is due to the short and rapid vegetative season of the Arctic system. Finally, it is worth considering the POC composition with depth to try to elucidate the origin and behavior of the particulate organic matter being exported. 3.8. Indications of the origin of particulate matter based on C/N and stable carbon isotope composition The exported POC may originate from primary or secondary autochtonous production or being released from ice-rafted debris. High export of ice-borne terrestrial organic matter in the Barents/ Fram Strait MIZ has been proposed (e.g., Wassmann et al., 2004). The bulk elemental and isotopic composition of the particles may provide some clues to this effect. First, station AO-01-1, located in Atlantic Water (Fig. 1) exhibit typical stable carbon isotope composition (d13C) of POC with values between –22 and –24% indicative of phytoplankton predominance throughout the top 100 m water column (Fig. 8A). This source is also supported by the C/N ratios in the range 6–7 (Fig. 8A). The C/N ratio in nearby ice and snow samples (collected with helicopter; see Gustafsson et al., 2005) was in range 28–48 indicative of terrestrial sources of this ice-rafted debris. However, it appears that the exported POC in the water column in this Atlantic Water location is not substantially affected by such an ice-rafted terrestrial signal. For a typical station in the FYI MIZ east of Svalbard in Arctic Water (station AO-01-2), the d13C POC depth profile starts around 30 in surface water and consistently increases with depth to – 27% at 100 m (Fig. 8B). We first note that due to Rayleigh fractionation, the carbon pools at these high latitudes are depleted by 1–2% relative to temperate zones. Secondly, it has been shown that POC sampled in the springtime under ice can be severely depleted relative to open water situations (e.g., Gustafsson et al., 2001). This may be a result of the regenerated inorganic carbon pool not being able to isotopically equilibrate with the d13C–CO2 in the overlying atmosphere. Instead, recycling of isotopically depleted carbon may drive down the d13C of the POC under the ice. We believe this is what the d13C POC data is reflecting at both the FYI station (Fig. 8B) and in the MYI conditions (Fig. 8C). Both those two ice-covered regimes start out with depleted d13C in the surface layer that consistently increase with several units with depth. These d13C POC depth profiles are reminiscent of the spring bloom profiles detailed for a boreal fjord (Waite et al., 2005). This feature may reflect settlinginduced isotopic fractionation whereby diatoms, known to be more depleted (Popp et al., 1998), are preferentially settling out of the surface ocean. Hence, despite the depleted d13C values, we believe the POC depth profile and settling flux are dominated by

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of 238U-234Th disequilibria in the northern Barents Sea MIZ and Nansen-Amundsen-Makarov basins during summer 2001 to demonstrate the importance of particle scavenging and settling export of organic matter to subsurface strata. The results show biogeochemically-consistent Arctic Ocean POC export fluxes with the small-volume direct at-sea beta counting method. The advent of this time and labor efficient approach opens the path for greater spatial coverage and resolution of POC export fluxes (see also e.g. Cai et al., 2010). Biogenic POC export in the Arctic Ocean clearly shows spatial and temporal variability, underlining the need for added observations. This study constrained the POC export fluxes in the 2001 season for northern Barents MIZ to 6– 18 mmol m  2 d  1 and for the high-latitude interior basins of the Eurasian Arctic to 2–3 mmol m  2 d  1. These values are consistent with most of the limited previous studies on POC export in these remote and poorly understood regimes. The now growing set of observations facilitates our ability to address what the Arctic Ocean ‘‘sink’’ of upper ocean POC export would be transformed to with anticipated Arctic change, including a forecasted 80% reduction in the summer ice pack (e.g., Johannessen et al., 2002). Clearly, our knowledge of these complex systems is incomplete. Nevertheless, there is an encouraging consistency between this and previous studies on the American side of the Arctic Ocean of high export efficiencies; 30–50% of primary production appears to be vertically exported as POC with no resolvable differences between shelves and interior systems. The melting ice would result in a widening of the MIZ. Earlier estimates by Reigstad and colleagues (2002) suggest that the putative new’’ MIZ-like’’ regimes would experience a tripling in new production for such interior Arctic Ocean areas. The differences in absolute export production fluxes constrained by the present and previous 234Th-based studies of POC fluxes between MIZ and Interior stations is on the order of 4–10 times. Whether it will turn out to be a factor of three or ten, extrapolated over 80% of the Arctic Ocean Proper, this would dramatically increase the sequestration of organic carbon in the Arctic Ocean. Expanded studies of POC export fluxes in the Eurasian Arctic sector, primarily the East Siberian Arctic Shelves and the Interior Basins, all within comprehensive biogeochemical frameworks, would be important contributions in the future.

100

Acknowledgments 120 Fig. 8. Upper ocean depth profiles of d13C POC at station AO-01-1 (MIZ, W. Atlantic Water; panel A), station AO-01-2 (MIZ, W. Arctic Water; panel B) and station AO-01-11 (High Arctic; Interior Arctic Water; panel C). The corresponding atomic C/N ratio at the same station in the same particles, mass-weight-integrated through the top 100 m, in ice and in snow, are inserted in the respective panel.

autotrophic production as opposed to terrestrial material released from the surrounding ice system. This is further supported by the C/N of the water column POC being in the range 6–10, whereas higher values were found in the ice-rafted terrestrial debris (Fig. 8B–C; Gustafsson et al., 2005). Taken together, this data supports that the exported POC is primarily of local biogenic origin.

We gratefully acknowledge the efforts of the crew of I/B ODEN, Ulf Hedman and other personnel at the Swedish Polar Research Secretariat, Paul Wassmann and Kalle Olli for kindly providing sediment trap splits for POC/Th determinations and for sharing ¨ ¨ and colleagues in the physical oceanotheir insights, Goran Bjork graphy group of AO-01 for sharing CTD data and providing Niskin samples, Don Porcelli for advising on Th chemistry, Zofia Kukulska for assistance in the laboratory and Torsten Persson for assistance with the beta counter. This study was financed by The Swedish Research Council (VR contract 1-AA/GB 12291-301). Constructive comments by three reviewers are appreciated. References

4. Implications and conclusions One of the key motivations behind expanding the database on upper ocean POC export fluxes in the Arctic Ocean is to allow us to make progress toward predictions of the consequences of ongoing Arctic climate change on the Arctic carbon cycle. The present study provides a large number of coherent measurements

Andreassen, I.J., Wassmann, P., 1998. Vertical flux of phytoplankton and particulate biogenic matter in the marginal ice zone of the Barents Sea in May 1993. Mar. Ecol. Prog. Ser. 170, 1–14. Baskaran, M., Swarzenski, P.W., Porcelli, D., 2003. Role of colloidal material in the removal of 234Th in the Canada Basin of the Arctic Ocean. Deep-Sea Res. I 50, 1353–1373. Bates, R.B., Hansell, D.A., Moran, S.B., Codispoti, L.A., 2005. Seasonal and spatial distribution of particulate organic matter (POM) in the Chukchi and Beaufort Seas. Deep-Sea Res. II 52, 3324–3343.

¨ . Gustafsson, P.S. Andersson / Deep-Sea Research I 68 (2012) 1–11 O

¨ Bjork, G., S¨oderkvist, J., Winsor, P., Nikolopoulos, A., Steele, M., 2002. Return of the cold halocline layer to the Amundsen Basin of the Arctic Ocean: Implications for the sea ice mass balance. Geophys. Res. Lett., 29, http://dx.doi.org/10.1029/2001GL014157. Buesseler, K.O., Bacon, M.P., Cochran, J.K., Livingston, H.D., 1992. Carbon and nitrogen export during the JGOFS North Atlantic Bloom Experiment estimated from 234Th: 238U disequilibria. Deep-Sea Res. I 39, 1115–1137. Buesseler, K.O., 1998. The decoupling of production and particulate export in the surface ocean. Global Biogeochem. Cycles 12, 297–310. Buesseler, K.O., Benitez-Nelson, C.R., Moran, S.B., Burd, A., Charette, M., Cochran, J.K., Coppola, L., Fisher, N.S., Fowler, S.W., Gardner, W.D., Guo, L.D., Gustafsson, ¨ ., Lamborg, C., Massque, P., Miquel, J.C., Passow, U., Santschi, P.H., Savoye, N., O Stewart, G., Trull, T., 2006. An assessment of particulate organic caron to thorium-234 ratios in the ocean and their impact on the application of 234Th as a POC flux proxy. Mar. Chem. 100, 213–233. Buesseler, K.O., Antia, A.N., Chen, M., Fowler, S.W., Gardner, W.D., Gustafsson, O., Harada, K., Michaels, A.F., Rutgers van der Loeff, M., Sarin, M.M., Steinberg, D.K., Trull, T., 2007. An assessment of the use of sediment traps for estimating upper ocean particle fluxes. J. Mar. Res. 65, 345–416. ¨ Cai, P., Rutgers van der Loeff, M., Stimac, I., Nothig, E.M., Lepore, K., Moran, S.B., 2010. Low export flux of particulate organic carbon in the central Arctic Ocean 234 238 as revealed by Th: U disequilibrium. J. Geophys. Res. 115, C10037. Carmack, E.C., 1986. Circulation and Mixing in Ice-Covered Waters. In: Untersteiner, N (Ed.), The Geophysics of Sea Ice. Plenum, New York, NY, USA, pp. 171–222. Charette, M.A., Moran, S.B., Pike, S, Smith, J.N., M., 2001. Investigating the carbon cycle in the Gulf of Maine using natural tracer thorium-234. J. Geophys. Res. 106, 11553–11579. Chen, M., Huang, Y., Cai, P., Guo, L., 2003. Particulate organic carbon export fluxes in the Canada Basin and Bering Sea as derived from 234Th/238U disequilibria. Arctic 56, 32–44. Chen, J.H., Edwards, R.L., Wasserburg, G.J., 1986. 238U,234U, and 232Th in seawater. Earth Planet Sci. Lett. 80, 241–251. Coppola, L., Roy-Barman, M., Wassmann, P., Mulsow, S., Jeandel, C., 2002. Calibration of sediment traps and particulate organic carbon export using 234 Th in the Barents Sea. Mar. Chem. 80, 11–26. Davis, J., Benner, R., 2005. Seasonal trends in the abundance, composition and bioavailability of particulate and dissolve organic matter in the Chukchi/ Beaufort Seas and western Canada Basin. Deep-Sea Res. II 52, 3396–3410. Dunne, J.P., Sarmiento, J., Gnanadesikan, A., 2007. A synthesis of global particle export from the surface ocean and cycling through the ocean interior and on the sea floor. Global Biogeochem. Cycles 21, GB4006, http://dx.doi.org/10.1029/ 2006GB002907. ¨ Fahl, K., Nothig, E.M., 2007. Lithogenic and biogenic particle fluxes on the Lomonosov Ridge (central Arctic Ocean) and their relevance for sediment accumulation: Vertical vs. lateral transport. Deep-Sea Res. Pt I: Oceanogr. Res. 54, 1256–1272. Gust, G., Bowles, W., Giordano, S., Huettel, M., 1996. Particle accumulation in a cylindrical sediment trap under laminar and turbulent steady flow: An experimental approach. Aquat. Sci. 58, 297–326. ¨ ., Haghseta, F., Chan, C., MacFarlane, J., Gschwend, P.M., 1997. Gustafsson, O Quantification of the dilute sedimentary ‘‘soot-phase: Implications for PAH speciation and bioavailability. Environ. Sci. Technol. 31, 203–209. ¨ ., Buesseler, K.O., Geyer, W.R., Moran, S.B., Gschwend, P.M., 1998. An Gustafsson, O assessment of the relative importance of horizontal and vertical transport of particle-reactive chemicals in the coastal ocean. Cont. Shelf Res. 18, 805–829. ¨ ., Nilsson, N., Bucheli, T.D., 2001. Dynamic colloid-water partitioning Gustafsson, O of pyrene through a coastal Baltic spring bloom. Environ. Sci. Technol. 35, 4001–4006. ¨ ., Andersson, P., Roos, P., Kukulska, Z., Broman, D., Larsson, U., Hajdu, Gustafsson, O S, Ingri, J., 2004. Evaluation of the collection efficiency of upper ocean subphotic-layer sediment traps: A 24-month in situ calibration in the open Baltic Sea using 234Th. Limnol. Oceanogr. Methods 2, 62–74. ¨ ., Andersson, P., Axelman, J., Bucheli, T.D., Komp, ¨ Gustafsson, O P., McLachlan, M.S., ¨ Sobek, A., Thorngren., J.O., 2005. Observations of the PCB distribution within and in-between ice, snow, ice-rafted debris, ice-interstitial water, and seawater in the Barents Sea marginal ice zone and the North Pole area. Sci. Total Envir. 342, 261–279. ¨ ., Larsson, J., Andersson, P., Ingri, J., 2006. The POC/234Th ratio of Gustafsson, O settling particles isolated using split flow-thin cell fractionation (SPLITT). Mar. Chem. 100, 314–322. Hargrave, B.T., von Bodungen, B., Stoffyn-Egli, P., Mudie, P.J., 1994. Seasonal variability in particle sedimentation under permanent ice cover in the Arctic Ocean. Cont. Shelf. Res. 14, 279–293. Johannessen, O.M., Bengtsson, L., Miles, M.W., Kuzmina, S.I., Semenov, V.A., Alekseev, G.V., Nagurny, A.P., Zakharov, V.F., Bobylev, L.P., Pettersson, L.H., Hasselmann, K., Cattle, H.P., 2002. Arctic Climate Change—Observed and Modeled Temperature and Sea Ice. NERSC Technical Report No. 218. Lalande, C., Lepore, K., Cooper, L.W., Grebmeier, J.M., Moran, S.B., 2007. Export fluxes of particulate organic carbon in the Chukchi Sea. A comparative study using 234Th/238U disequilibria and drifting sediment traps. Mar. Chem. 103, 185–196. Lalande, C., Moran, S.B., Wassmann, P., Grebmeier, J.M., Cooper, L.W., 2008. 234Thderived particulate organic carbon fluxes in the northern Barents Sea with comparison to drifting sediment trap fluxes. J. Marine Syst. 73, 103–113. Lalande, C., Be´langer, S., Fortier, L., 2009. Impact of a decreasing sea ice cover on the vertical export of particulate organic carbon in the northern Laptev Sea,

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Siberian Arctic Ocean. Geophys. Res. Lett. 36, L21604–L21610, doi:10.21029/ 22009GL040570. Lepore, K., Moran, S.B., 2007. Seasonal changes in thorium scavenging and particle aggregation in the western Arctic Ocean. Deep-Sea Res. I 54, 919–938. Moran, S.B., Buesseler, K.O., 1993. Size-fractionated 234Th in continental shelf waters off New England: implications for the role of colloids in oceanic trace metal scavenging. J. Mar. Res. 51, 893–922. Moran, S.B., Ellis, K.M., Smith, J.N., 1997. 234Th/238U disequilibria in the central Arctic Ocean: implications for particulate organic carbon export. Deep-Sea Res. II 44, 1593–1606. Moran, S.B., Smith, J.N., 2000. 234Th as a tracer of scavenging and particle export in the Beaufort Sea. Cont. Shelf Res. 20, 153–167. Moran, S.B., Weinstein, S.E., Edmonds, H.N., Smith, J.N., Kelly, R.P., Pilson, M.E.Q., Harrison, W.G., 2003. Does 234Th/238U disequilibrium provide an accurate record of the export flux of particulate organic carbon from the upper ocean? Limnol. Oceanogr. 48, 1018–1029. Moran, S.B., 2004. Regional variability in POC export flux in the Arctic Ocean determined using 234Th as a tracer. In: Stein, R., Macdonald, R.W. (Eds.), The Organic Carbon Cycle in the Arctic Ocean.. Springer, Berlin, Ch. 5.4. Moran, S.B., Kelly, R.P., Hagstrom, K., Smith, J.N., Grebmeier, J.M., Cooper, L.W., Cota, G.F., Walsh, J.J., Bates, N.R., Hansell, D.A., Maslowski, W., Nelson, R.P., Mulsow, S., 2005. Seasonal changes in POC export flux in the Chukchi Sea and implications for water column-benthic coupling in Arctic shelves. Deep-Sea Res. II 52, 3427–3451. Olli, K., Wexels Riser, C., Wassmann, P., Ratkova, T., Arashkevich, E., Pasternak, A., 2002. Seasonal variation in vertical flux of biogenic matter in the marginal ice zone and the central Barents Sea. J. Marine Syst. 38, 189–204. Olli, K., Wassmann, P., Reigstad, M., Ratkova, T.N., Arashkevich, E., Pasternak, A., Matrai, P.A., Knulst, J., Tranvik, L., Klais, R., Jacobsen, A., 2007. The fate of production in the central Arctic Ocean—top-down regulation by zooplankton expatriates? Results from a 3 week ice drift at 881N. Prog. Oceanogr. 72, 84–113. Popp, B.N., Laws, E.A., Bidigare, R.R., Dore, J.E., Hanson, K.L., Wakeham, S.G., 1998. Effect of phytoplankton cell geometry on carbon isotopic fractionation. Geochim. Cosmochim. Acta 62, 69–77. ¨ ygarden, S., Rey, F., 2002. Variations in Reigstad, M., Wassmann, P., Riser, C.W., O hydrography, nutrients, and chlorophyll a in the marginal ice-zone and the central Barents Sea. J. Marine Syst. 38, 9–30. Reigstad, M., Carroll, J., Slagstad, D., Ellingsen, I., Wassmann, P., 2011. Intraregional comparison of productivity, carbon flux and ecosystem composition within the northern Barents Sea. In: Arctic Marine Ecosystems in an Era of Rapid Climate Change, special issue, Progr. Oceanogr., vol. 90, pp. 33–46. Rutgers van der Loeff, M.M., Moore, W.S., 1999. Determination of natural radioactive tracers. In: Grasshoff, K., Ehrhardt, M., Kremling, K. (Eds.), Methods of Seawater Analysis. Verlag Chemie, Weinheim, Germany, pp. 365–397, Chapter 13. Slagstad, D., Ellingsen, I.H., Wassmann, P., 2011. Evaluating primary and secondary production in an Arctic Ocean viod of summer sea ice: An experimental simulation approach. In: Arctic Marine Ecosystems in an Era of Rapid Climate Change, special issue, Progr. Oceanogr., vol. 90, pp. 117–131. ¨ ., 2004. Latitudinal fractionation of polychlorinated Sobek, A, Gustafsson, O biphenyls in surface seawater along a 621N–891N transect from the southern Norwegian Sea to the North Pole area. Environ. Sci. Technol. 38, 2746–2751. ¨ ., 2006a. Partitioning of polychlorinated Sobek, A, Reigstad, M., Gustafsson, O biphenyls between Arctic seawater and size-fractionated zooplankton. Environ. Toxicol. Chem. 25, 1720–1728. ¨ ., 2006b. On the relative significance of bacteria for Sobek, A., Olli, K., Gustafsson, O the distribution of polychlorinated biphenyls in Arctic Ocean surface waters. Environ. Sci. Technol. 40, 2586–2593. Steinberg, D.K., Pilskaln, C.H., Silver, M.W., 1998. Contribution of zooplankton associated with detritus to sediment trap ‘‘swimmer’’ carbon in Monterey Bay, CA. Mar. Ecol. Prog. Ser. 164, 157–166. Trimble., S.M., Baskaran, M., 2005. The role of suspended particulate matter in 234 Th scavenging and 234Th-derived export fluxes of POC in the Canada Basin of the Arctic Ocean. Mar. Chem. 96, 1–19. ¨ ., Lindahl, O., Tiselius, P., 2005. Linking ecosystem Waite, A., Gustafsson, O dynamics and biogeochemistry: Sinking fractionation of organic carbon in a Swedish fjord. Limnol. Oceanogr. 50, 658–671. Wassmann, P., 2002. Seasonal C-cycling variability in the open and ice-covered waters of the Barents Sea: an introduction. J. Marine Syst. 38, 1–7. Wassmann, P., Bauerfeind, E., Fortier, M., Fukuchi, M., Hargrave, B., Moran, B., Noji, ¨ T., Nothig, E.-M., Olli, K., Sasaki, H., Shevchenko, V., 2004. Particulate organic carbon flux to the Arctic Ocean sea floor. In: Stein, R., Macdonald, R.W. (Eds.), The Organic Carbon Cycle in the Arctic Ocean. Springer, Berlin, Ch.5. Wassmann, P., 2004. Particulate organic carbon flux to the seafloor of the Arctic Ocean: Quantity, seasonality and processes. In: Stein, R., Macdonald, R.W. (Eds.), The Organic Carbon Cycle in the Arctic Ocean. Springer, Berlin, Ch. 5.5. Wassmann, P., Olli, K., 2004. Central Barents Sea and Northern Spitsbergen. In: Stein, R., Macdonald, R.W. (Eds.), The Organic Carbon Cycle in the Arctic Ocean. Springer, Berlin, Ch. 5.3.4. Wassmann, P., special issue editor, 2011. In: Arctic Marine Ecosystems in an Era of Rapid Climate Change, special issue, Progr. Oceanogr, vol. 90, pp. 1–17. Wheeler, P.A., Watkins, J.M., Hansing, R.L., 1997. Nutrients, organic carbon and organic nitrogen in the upper water column of the Arctic Ocean: implications for the sources of dissolved organic carbon. Deep-Sea Res. II 44, 1571–1592.