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JOURNAL OF GEOPHYSICAL RESEARCH: OCEANS, VOL. 118, 2133–2154, doi:10.1002/jgrc.20101, 2013

Canadian Basin freshwater sources and changes: Results from the 2005 Arctic Ocean Section Robert Newton,1 Peter Schlosser,1,2 Richard Mortlock,3 James Swift,4 and Robie MacDonald5 Received 11 October 2012; revised 14 January 2013; accepted 27 January 2013; published 29 April 2013.

[1] We present measurements of oxygen isotope ratios and nutrient concentrations along

the 2005 Arctic Ocean Section aboard the icebreaker Oden. The data are used to estimate freshwater contributions from meteoric water (mainly river runoff), sea-ice meltwater, and Chukchi Sea shelf water, itself a combination of Pacific and indigenous Arctic water types. Nutrients ratios are combined to form quasi-conservative water-mass tracers (phosphate-star, N-star, and the empirical Arctic N-P relationship) and used along with salinity and d18O, which are conservative in the ocean interior. Disagreements between two different freshwater analyses in the Western Arctic are largely resolved using a salinity-dependent Redfield ratio, a new estimate of the Pacific end-member, and an analysis of the Bering Strait inflow contribution to detraining shelf waters. Freshwater components from 2005 are placed into the context of the overlapping 1994 Arctic Ocean Section (aboard the Louis St. Laurent) and a time series of hydrographic/tracer casts between 1987 and 1992 in the Canada Basin. Compared to 1987–1994; the 2005 transect exhibits increased meteoric water concentrations in the northern part of the Canadian Basin and a decrease in the southern part. This pattern is related to changes in the distribution of wind-stress curl during the several years prior to each sampling campaign. In addition, a previously observed correlation between sea-ice formation and river runoff disappears over the Central Arctic in 2005, a change that we attribute to a northward shift of sea-ice formation. Resampling approximately every 3 years should resolve the dynamics driving changes in freshwater and nutrient distributions. Citation: Newton, R., P. Schlosser, R. Mortlock, J. Swift, and R. MacDonald (2013), Canadian Basin freshwater sources and changes: Results from the 2005 Arctic Ocean Section, J. Geophys. Res. Oceans, 118, 2133–2154, doi:10.1002/jgrc.20101.

1.

Introduction

[2] In the Arctic Ocean and the subarctic seas, the distribution of salinity exerts primary control over density gradients and thus plays a major dynamical role. In the vertical, salinity gradients account for nearly all of the stability of the Arctic water column. The flux of heat from the relatively warm Atlantic Layer to the base of the Arctic sea-ice cover and the flux of nutrients to the photic zone are both strongly modulated by changes in this stratification. Erosion of the freshwater anomaly beneath the sea-ice has, for example, been implicated as a contributing factor in the preconditioning of the perennial Arctic sea ice cover for its recent rapid decline [see Polyakov et al., 2010, and references therein]. Horizontal salinity differences, integrated through the water column, dominate steric height gradients, 1

Lamont-Doherty Earth Observatory, Palisades, New York, USA. Department of Earth and Environmental Sciences and Department of Earth and Environmental Engineering, Columbia University, New York, New York, USA. 3 Rutgers University, Newark, New York, USA. 4 University of California, San Diego, California, USA. 5 Institute of Ocean Sciences, Department of Fisheries and Oceans, Canada. 2

Corresponding author: R. Newton, Lamont-Doherty Earth Observatory, Palisades, NY 10964-8000, USA. ([email protected]) ©2013. American Geophysical Union. All Rights Reserved. 2169-9275/13/10.1002/jgrc.20101

which are important in establishing the Arctic Ocean’s quasi-permanent circulation features. These include coastal currents that carry European river runoff eastward as far as the Bering Strait, the Trans-Polar Drift (TPD) that reaches from Siberia to Fram Strait, the Beaufort Gyre (BG), and the East Greenland Current (EGC). The latter is a primary source of buoyancy, in the form of a negative salinity anomaly, to the surface of the Greenland and Iceland seas. These are critical sites for the convective formation of deep waters that eventually exit the Nordic seas across the Scotland/Iceland/Greenland sill system and participate in the formation of North Atlantic Deep Water (NADW) [e.g., Schlosser et al., 1991; Boenisch and Schlosser, 1995; Mauritzen and Hakkinen, 1999, and references therein]. [3] Temporal changes in the freshwater content of the Arctic Ocean and its peripheral seas are reviewed by Polyakov et al. [2008], on the basis of the past 100 years of measurements of salinity and temperature. On both inter-annual and multi-decadal timescales, the Central Arctic undergoes significant swings in its freshwater content. Given the magnitude of changes in freshwater exchange with the atmosphere (which are small by comparison), the observations imply large changes in the flux from the Arctic to the sub-Arctic, as well as changes in the direction of flux between the Arctic and its peripheral shelf seas. Inter-annual fluxes on the order of 2000 km3/yr and sustained decadescale anomalies of about 200 km3/yr are common in the

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record [Polyakov et al., 2008; Serrezze et al., 2006]. The periodic flushing of anomalously large volumes of sea ice and freshwater from the central Arctic through Fram Strait has been implicated in the modulation of stratification in the Nordic seas and the North Atlantic and related deep convection rates, a basic climate variable [Dickson et al., 1988; Mauritzen and Hakkinen, 1997; Karstensen et al., 2005; Boenisch and Schlosser, 1995; Holfort et al., 2008]. Salinity measurements, from which Polyakov et al. [2008] construct freshwater inventories and fluxes, are insufficient to illuminate the individual freshwater sources and their contributions to both inventories and fluxes. A central goal of stable isotope ratio analysis in Arctic and sub-Arctic waters is to understand the underlying dynamics of freshwater component distribution and variations in the Arctic. [4] Most of the water in the Arctic Ocean is derived from North Atlantic sources, with a relatively constant salinity between about 34.9 and 34.95 psu. The large salinity gradients within the Arctic Ocean and its shelf seas result from the addition of freshwater from non-Atlantic sources: meteoric water including river runoff and local precipitation or evaporation, Pacific Water which is approximately 2.5 psu fresher than inflowing Atlantic Water, and formation and melting of sea ice. The detailed modeling of freshwater distributions, sources, and sinks in the Arctic Ocean has therefore been a focus of several studies in recent years, with significant progress being made during the last decade. Recent model studies have, for example, been able to replicate the overall distribution of freshwater in the Arctic, decadal shifts of freshwater pathways linked to changes in the large-scale wind patterns, and the periodic pulses into the Nordic Seas known as “Great Salinity Anomalies” [see, for example, Jahn et al., 2010; Newton et al., 2006; Newton et al., 2008; Karcher et al., 2012, and references therein].

[5] To be quantitatively useful, models must be validated against observations, and the sources of freshwater and brine in the Arctic are not directly observable. Several proxies are useful for decomposing Arctic Ocean water with regard to the sources of its freshwater composition and its anomalies. In this study, we use the stable isotope ratio of water, 16 H18 2 O/H2 O, which is much lower in high-latitude precipitation than in sea-ice meltwater or ocean water, together with combinations of nutrients that can distinguish between Pacific and Atlantic inflows. [6] In addition to reporting water mass decompositions from the Arctic Ocean Section 2005 (AOS05) and comparing them with earlier data, we consider some technical aspects of water mass decomposition in the Arctic. The choice of proxy for the Pacific inflow is complicated by the fact that unlike isotope ratios in water, nutrients are not conservative tracers. We compare different methods of combining nutrients in an attempt to account for biological processes. We also consider a practice that has become common in analyzing Arctic water samples for freshwater sources: first splitting the sample into “fresh” and “salty” components and then independently decomposing the fresh component to meteoric and seaice contributions and the saline component into Atlantic and Pacific contributions.

2.

Sample Collection and Measurement

[7] The central data set presented here was collected during the 2005 Arctic Ocean Section from the Alaskan continental slope to the Nansen Basin (Figure 1). We compare the 2005 results with data from earlier campaigns including the 1994 Arctic Ocean Section (AOS94) that passed close to the 2005 track near the North Pole and in the Makarov

Alaska

Lo m

on

os

ov R

idg

e

Canadian Basin

Canadian Arctic Archipelago

Greenland

Figure 1. Hydrographic stations: 2005: magenta, 1994: green, 1987-1991: yellow. Circled stations were used in the comparison of profiles in the Makarov Basin. Makarov Basin stations are labeled in the inset. 2134


Basin, the 1997/98 SHEBA encampment drift that passed near the 2005 track on the northern flank of the Chukchi Cap, and a site in the southern Canada Basin, at about 72.7N, 143.4W, that the Institute for Ocean Studies (IOS) repeatedly sampled, which they refer to as “Station A.” [8] The AOS05 trans-Arctic section was completed between 19 August and 25 September 2005 aboard the Swedish icebreaker ODEN (Figure 1). The section begins north of Barrow, Alaska, at 71.38N, 152.29W, in approximately 100 m water depth over the Alaskan continental shelf, and ends off the Barents Sea shelf, at 81.85N, 33.56E. The cruise proceeded directly North from Barrow, trended northwestward across the central Makarov Basin to the Lomonosov Ridge, and continued to the North Pole along the ridge. At Stations 6 and 7 in the Canada Basin, technical problems prevented the collection of water in the upper 500 m of the water column resulting in a data gap across part of the Beaufort Gyre. Station coverage in the Nansen Basin was very sparse due to time constraints. Thus, we focus our analysis on the section between Station 3 in the southern Canada Basin and Station 50 near the southern boundary of the Amundsen Basin. Along that track, the cruise offers the most detailed scientific Arctic Ocean transect to date [see also Bjork et al., 2007]. [9] At each of the 53 hydrographic stations depth profiles for salinity, temperature, dissolved oxygen, fluorescence, and nutrients (silicate, phosphate, nitrate, and nitrite) were completed. At most stations, micro-structure casts were performed, and between stations expendable CTD profilers (XCTDs) were deployed to measure temperature and salinity in approximately the upper 1000 m. Tracers including Chlorofluorocarbons (CFCs), oxygen isotopes, helium isotopes, and tritium were measured at 42 of the 53 stations. At a smaller subset of stations, additional tracers were collected from Niskin bottles: colored dissolved organic matter (CDOM: refractory dissolved organic molecules), “triple” oxygen isotopes, 14C, and 99Tc (technetium). Temperature and fluorescence were measured by CTD sensors. Salinity was measured both by CTD and bottle measurements. [10] All water samples were taken from 10 L Niskin bottles arranged in a 36-position rosette, with a multi-probe sonde at its center that housed sensors for conductivity, pressure, dissolved oxygen, and fluorescence. Here we use the salinity and dissolved oxygen data from the bottle samples. Water samples for measurement of 18O/16O ratios of water were collected in 50 ml glass bottles. The bottles were triple rinsed in water from the Niskin bottle to be sampled, filled, and sealed using polypro-lined caps and electrical tape. 16 [11] Oxygen isotope ratios of water (H18 2 O/H2 O) were measured at Rutgers University using the method described by Fairbanks [1982] and are reported as d18O, the per mil anomaly relative to Vienna Standard Mean Ocean Water (VSMOW). Analytic precision (1 sigma) was approximately +/0.025%. The data were quality controlled by comparing the individual measurements to a smoothed approximation to the d18O profile at each station. A local second- or thirdorder polynomial was used to fit the data in a least-squares mode, and residuals from the data to the profile were used to identify possible fliers as well as to evaluate the actual scatter in the data set. Scatter around the profiles was mostly

within the analytical error, although for a few profiles, the root-mean-square residual around the smoothed profiles was slightly larger. [12] Nutrient concentrations, required for the identification of Pacific and Arctic shelf waters, were measured aboard the Icebreaker ODEN in accordance with established WOCE (World Ocean Circulation Experiment) standard techniques [Gordon et al., 1992]. Measurements were made using an Ocean Data Facility (ODF)-modified 4-channel Technicon AutoAnalyzer II, generally within 2 h of sample collection. Dissolved oxygen concentrations were measured onboard ODEN using an ODF-designed automated oxygen titrator based on the Winkler titration technique described by Carpenter [1965]. [13] It needs to be noted that the presence of a large icebreaker and the insertion of a rosette that is nearly 2 m in diameter into the water disturbs the water’s surface. The upper approximately 10 m of the water column is more or less thoroughly mixed, and the air-sea gas exchange is increased dramatically. In the following, we avoid drawing conclusions based on the vertical structure within the upper 20 m of the water column. Any structure in this depth range has likely been altered by the sampling process itself. Fortunately, most measurements used in this study (oxygen isotope ratios, nutrients, salinity, and potential temperature) are not very sensitive, on the timescale of the sampling process, to increased gas exchange. The greatest impact is likely in the oxygen concentration, for which increased stirring is likely to increase the degree and depth of equilibration. Thus, despite the vertical mixing at the surface, changes in the vertically integrated property inventories between stations should be valid.

3.

Hydrographic Setting

[14] Relatively warm (>0 C), salty (>34.9 psu) water flows into the Arctic along the European slope (Figures 2a and 2b). It passes around the Arctic Basin from West to East in an advective boundary current, hugging the continental slope, that slowly cools, freshens, and deepens along its path. This boundary current has two main cores: the Fram Strait Branch Water (FSBW) enters the Arctic through Fram Strait and flows eastward along the continental slope [Schauer et al., 2002]. The Barents Sea Branch Water (BSBW) enters the Arctic as a broad current through Bear Trough, between Svalbard and Norway, crosses the Barents Sea and enters the boundary current in the vicinity of Saint Anna Trough [Rudels et al., 2000; Schauer et al., 2002; Mauldin et al., 2010]. Modified by a wide range of conditions at the surface in the Barents Sea, BSBW spans a wider range of densities than FSBW and fills a correspondingly wider depth range as it detrains from the shelf North and East of the Barents Sea. Atlantic Water from the boundary current mixes into the interior along isopycnal surfaces, covering the basin between about 200 and 600 m depth with an “Atlantic” layer that is easily distinguished as a vertical temperature maximum throughout the Arctic basin. [15] At the surface is a thin layer of relatively fresh “Polar Water” which is buoyant as a result of inputs from river runoff, local precipitation, and melting sea ice. Between the Atlantic and Polar waters is a complex, strongly stratified halocline. There may be one, two, or three distinct vertical

2135


a

4

8

12

Salinity (psu) 20 30

50

4

52

8

Nitrate (μmol/kg) 12 20 30

50

52

f

Phosphate (μmol/kg)

Temperature

b

g

Arctic N/P Tracer (fraction)

Freshwater Anomaly (percent, S0 = 34.92)

h

c

Phosphate Star (PO4*) (μmol/kg)

Oxygen Isotope Ratios (δ18O, per mil)

d

i

Dissolved Oxygen Anomaly (fraction)

Oxygen (μmol/kg)

e

j

Major Basins 1000 Canada 2000 3000 4000

1000 2000 3000 4000

Nansen Canada Makarov Amundsen

Major Ridge Systems North Wind Alpha Lomonsov Gakkel

Figure 2. Data and combined tracers along the upper 500 m of the AOS 2005 cruise track. layers within the halocline, depending on location and season [Aagaard et al., 1981; Steele et al., 1995; Steele and Boyd, 1998; Martinson and Steele, 2001; Boyd et al., 2002]. At its base is a layer that mixes with the top of the Atlantic Layer and is generally known as the Lower Halocline Layer (LHL), between salinities of about 34.1 and 34.8 [Rudels et al., 1996]. Over much of the Canadian Basin, and sometimes extending into the Amundsen Basin as well, the so-called Upper Halocline Layer (UHL), centered around a salinity of about 33.1 [Jones and Anderson, 1986] and a depth of about 150 m [Moore et al., 1983; Schlosser et al., 2002], is observed. In the Canadian Basin, the UHL can be identified readily by a strong peak in nutrient concentrations. Over some parts of the Arctic, there is a Cold Halocline immediately below the mixed layer which is close to the freezing point throughout its vertical extent, but strongly stratified in salinity.

[16] The halocline is fed both from above, by vertical mixing and horizontally, by detrainment from the continental shelves [Aagaard et al., 1981; Steele et al., 1995]. Vertical mixing is largely a seasonal process, with sea ice meltwater and river runoff stratifying the water column in the warm months and sea-ice formation adding dense brine that drives convection during fall and winter. Detrainment is an episodic, highly variable process that involves interactions between atmospheric forcing and oceanic currents over the broad Arctic shelves [Bauch et al., 2009; Ekwurzel et al., 2001]. Plumes of relatively salty, brine-enriched water, descending along the continental slope, spread northward along isopycnal surfaces over the deep central Arctic. Over the deep Arctic basins, the shelf-influenced waters have much more consistent properties than over the shelves themselves, a sign of strong isopycnal mixing after detrainment from the shelf.

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[17] The horizontal gradients of wind stress play a major role in organizing the distribution of freshwater in the Arctic surface layer [Hunkins and Whitehead, 1992; Proshutinsky and Johnson, 1997; Proshutinsky et al., 2002, 2009; Newton et al., 2006, 2008; Jahn et al., 2010; Bauch et al., 2011; Alkire et al., 2007]. On average, high pressure dominates over the Canadian Basin and low pressure over the Nansen Basin, which leads to (1) the TPD from Siberia to Greenland and (2) Ekman convergence in the BG and divergence on the Eurasian side of the TPD. One result is a downward bowing of isopycnals and a relatively thick cap of fresh Polar Water in the BG, approximately between stations 5 and 25 on the AOS05 cruise track. This general hydrographic structure of the water column is characterized by significant shifts of the distribution of freshwater on timescales of years to decades [see, e.g., Newton et al., 2006, 2008; Polyakov et al., 2008; Jahn et al., 2010; Morison et al., 1998; Bauch et al., 2009; Anderson et al., 2004; Proshutinsky and Johnson, 1997; Morison et al., 2012, and references therein]. These shifts can also occur on seasonal timescales [Bauch et al., 2011]. In order to understand these shifts, it is useful to decompose the freshwater signal into its components: meteoric water, precipitation and evaporation, sea-ice meltwater, and the freshwater component of Pacific inflow. Salinity distinguishes the Atlantic fraction of a sample from the components that add a freshwater signal. Oxygen isotope ratios of water distinguish sea-ice meltwater from meteoric water (river runoff plus local precipitation/ evaporation). Nutrient concentrations can be used to distinguish between Pacific inflow (which has relatively high nutrient values) and Atlantic inflow (relatively low nutrient concentrations). [18] Salinity and oxygen isotope ratios are set at the boundaries of the Arctic Ocean. Within the Arctic Ocean interior, they are conservative. [19] Nutrients are subject to biological cycling, mainly via photosynthesis, respiration, nitrogen fixation, and denitrification. To minimize the impact of photosynthesis and respiration, the two dominant processes leading to non-conservative behavior, typically combinations of nutrients that are quasiconservative are calculated (for details, see below).

4.

Results

[20] We measured 549 d18O samples taken in the upper 500 m of the water column along the AOS05 cruise track (Figure 2d). The lowest d18O values (between about 1.5 and 3.5 per mil), indicative of meteoric water, are observed in the upper approximately 50 m of the water column from the Canadian continental shelf to the Gakkel Ridge, which separates the Amundsen and Nansen basins. This pattern is roughly congruent with the pattern of low-salinity anomalies (Figure 2a), indicating qualitatively that meteoric water is the primary factor in forming Polar water in the upper water column. Atlantic water is visible flowing into the Arctic as a thick layer with positive d18O values (approximately 0.2 to 0.3 per mil) and high salinities (>34.9) on the European side. From about 400 m downward, positive d18O values extend across the Arctic, although there is a subtle tendency towards lower d18O values on the Canadian side of the section, indicating dilution of the Atlantic inflow by plumes of Arctic and/or Pacific origin detraining from the shelves.

[21] In near-surface waters, well-defined plumes of low d18O waters (d18O < 2.5) are evident over the Lomonosov Ridge (LR), the Alpha-Mendeleyev Ridge system (AMR), and in the BG. Anti-cyclonic wind stresses dominate, on average, over the Canadian Basin, and the wind-stress curl gradient between the Canadian and Eurasian basins drives Ekman transport into the BG, forming a reservoir of low density water with high meteoric water content. [22] The high river runoff signals over the LR and AMR result from topographically trapped components of the general drift from Siberia toward Fram Straight. Steric height gradients of the order of 0.5 m between the Siberian and Alaskan shelves and the North Atlantic, together with the average wind stress patterns between the Canadian and Eurasian basins, drive a flow from the shelves toward Fram Strait and into the Nordic seas [e.g.; Newton et al., 2008, and references therein]. Water from the continental shelves preferentially detrains at saddle features where the major ridges join the continental shelf [Newton et al., 2008]. Less energy input is required at these features to balance vorticity changes as the thickness of isopycnal layers increases. Thus, the trans-Arctic transport tends to follow the flanks of the ridges as it crosses the central basin, [Holloway and Wang, 2009; Newton et al., 2008]. The low-d18O surface plumes centered at stations 25 (over the AMR) and between stations 32 and 43 (over the LR) sample these topographically trapped plumes of runoff-rich shelf water. [23] The dominant feature of the phosphate distribution in the Arctic Ocean is a tongue of high concentrations extending from the North American continental shelf northward across the Canadian Basin (Figure 2g). High phosphate concentrations are centered on a salinity horizon of 33.1, which sits at a depth of about 150 m in the southern Canada Basin and shoals upward to the North. At the shelf break, the layer is nearly 200 m thick, but thins as it shoals northward, and disappears in the northern part of the Canadian Basin. Vertical phosphate maxima at the stations that sample this feature are between approximately 1.6 and 2.05 mmol kg1. A secondary peak of phosphate, about 0.9 mmol kg1, shows up over the LR, coinciding with the runoff-rich plume evident in the d18O distribution. [24] Even where it is lowest, in surface waters over the northern Makarov Basin and over the Gakkel Ridge, the phosphate concentration is never insignificant. The lowest phosphate concentration from the ODEN cruise (1106 good nutrient data points) is 0.23 mmol kg1. Nitrate concentrations, on the other hand, approach zero in the surface waters of the entire Canadian Basin (Figure 2f). Thus, it appears that primary production in the pelagic Arctic is nitrate limited. In the photic zone, horizontal structure is visible in the phosphate field that is absent in nitrate (which does not exceed about 1 mmol kg1 in this zone). Below the photic zone, the patterns of the phosphate and nitrate fields are remarkably similar. Both are dominated by the intense tongue of shelf-derived water in the Canadian Basin, below which they decline gradually with depth. Below 1000 m depth (not shown), the nitrate and phosphate concentrations slowly increase with depth. In the abyss, nitrate is between about 14.5 and 15 mmol kg1 and phosphate is between 1.0 and 1.05 mmol kg1. [25] Outside the high-nutrient tongue, down to about 1000 m, the nitrate and phosphate fields exhibit a more

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subtle horizontal heterogeneity. There is a minimum centered at station 9, under the Beaufort Gyre, a peak at station 21, near the AM-Ridge, a sharp gradient from higher to lower concentrations as the section passes over the LR into the Amundsen Basin, and then another increase centered over the Gakkel Ridge. These shifts are small, i.e., less than 0.1 mmol kg1 for phosphate and less than 1 mmol kg1 for nitrate. However, they are well correlated (r = 0.87, p < 103) for phosphate and nitrate, with an empirical Redfield ratio of 14.6 and a phosphate excess (nitrate = zero intercept on an N/P plot) of 0.22 mmol kg1. We compared the nutrient concentrations between 200 and 1000 m with transient tracer ages (not reported here). Those results indicate that nutrient peaks in this band represent older waters (in the sense that they have been isolated from the mixed layer for a longer time). Thus, it is likely that these lateral gradients are caused by circulation and spreading patterns and the related gradients in the mean ventilation times of halocline waters. [26] The dissolved oxygen (DO) distribution (Figure 2j) is in most respects complementary to the nitrate and phosphate fields (Figures 2f and 2g). In the upper 25 m, nitrate concentrations have been drawn down by productivity to below 1 mmol kg1 and oxygen concentrations are set by exchange with the atmosphere to values approximately 3.75% above saturation. As nitrate and phosphate concentrations increase with depth, oxygen concentrations decrease. Horizontally, dissolved oxygen shows the same subtle gradients as the nutrient fields but with an opposite sign. Assuming that these gradients represent nutrient regeneration, the empirical Redfield ratio for dissolved oxygen to nitrate (8.2) in the upper 500 m is lower than global averages.

5.

Discussion

5.1. Freshwater Sources [27] Our goals are to understand the role played by freshwater sources in the large-scale Arctic Ocean circulation, and to expose the mechanisms that drive changes in the circulation, as well as the evolving interaction between the Arctic Ocean and the regional climate. We focus on freshwater because, as described in the introduction, it is the addition of low-salinity water to the Arctic surface that dominates the distribution of buoyancy in the Arctic Ocean. In principle, as long as water masses can be cleanly separated in an appropriate parameter space, one can separate a water sample into its component parts. In the case of the Arctic Ocean, exchanges with the North Atlantic are one to two orders of magnitude larger than exchanges with the North Pacific, the atmosphere, or the surrounding land masses. Thus, we take the point of view that the basic water mass is North Atlantic water, to which Pacific Inflow, precipitation and river runoff have been added. Since North Pacific water is significantly fresher than North Atlantic water (S ~ 32.5 versus S ~ 34.9); all three of these additions are freshwater sources. In addition, the annual sea-ice melt/freeze cycle separates the surface water into nearly fresh sea ice (S ~ 4) and salty brines. [28] We need a parameter space in which these sources are well separated. We do not have a tracer that separates precipitation from river runoff, so we group those two sources together as “meteoric water.” Salinity strongly separates the meteoric water (S ~ 0) and sea-ice meltwater

(S ~ 4) from the saline waters: North Atlantic (S ~ 34.9) and North Pacific inflow (S ~ 32.5). Meteoric water can be separated from the other water masses by oxygen iso16 tope ratios. Oxygen isotope ratios in water (H18 2 O/H2 O) shift due to fractionation during phase changes: by approximately 9 per mil during evaporation or condensation and about 2.6 per mil during freezing [Ekwurzel et al., 2001]. As a result, precipitation at high latitudes has a significant 18 O deficit, from repeated cycles of evaporation and precipitation (Rayleigh distillation). Typical average Arctic river runoff and precipitation d18O values are about 20 per mil [Oestlund and Hut, 1984; Bauch et al., 1995; Ekwurzel et al., 2001] compared to approximately +0.3 per mil for North Atlantic water, and sea-ice values are about 2 to 2.6 per mil above the water from which the ice has been formed. In the ocean interior, d18O is only impacted by water-mass mixing and has an excellent dynamic range for differentiating meteoric water from sea-ice meltwater. [29] North Pacific water is richer in nutrients than North Atlantic water. The North Atlantic drift, which is the source for Atlantic inflow to the Arctic, is composed of water that has been at the surface for a long time and has, therefore, been depleted in its nutrients. Nutrients in North Pacific water, on the other hand, are replenished by upwelling waters that have been isolated from the atmosphere for a long time and thus are repleted with inorganic nutrients. Therefore, we can use nutrient concentrations to distinguish between Pacific water inflowing through Bering Strait and Atlantic water, inflowing between Greenland and Norway. (This simple picture is complicated somewhat by Arctic shelf processes, which will be discussed in some detail below.) [30] Sea-ice meltwater is distinguished from saline waters by its low salinity and from river runoff by its high d18O values. However, establishing accurate values for the assumed sea-ice meltwater mass, or “end-member”, is complicated by the fact that the ice is mobile and can undergo multiple formation/melt cycles along its pathway. During its formation, sea ice rejects of the order of 80% of all dissolved ions and, as noted above, fractionates oxygen isotopes by about +2.6 per mil. In our analyses, we assumed that the sea-ice end-member d18O value for each sample is 2.6 per mil above surface water values at each profile. This is equivalent to assuming that the sea-ice and mixed layer move together and that in the upper 500 m, the sea-ice meltwater in samples is related to the vertical convective flux more than it is to the horizontal mixing along isopycnals. These assumptions are not strictly accurate, but we have found them to be a workable compromise [Ekwurzel et al., 2001, and references therein; MacDonald et al., 2002]. 5.2. Freshwater Component Separation 5.2.1. Method [31] The basic method for water-mass decomposition is straightforward. If a water sample can be assumed to be a combination of several water mass types, each with distinct and consistent properties, then for each measured property we can write down a linear equation: f1 ½x1 þ f2 ½x2 þ f3 ½x3 þ ⋯ ¼ ½x

(1)

where fi is the component fraction and [x]i is the concentration of property x in the ith water mass, or end-member. If

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a

three parameters are measured (e.g., salinity, d18O, and a nutrient, N), and the sample is assumed to be composed of four end members (e.g., Atlantic, Pacific, meteoric and sea-ice melt waters), then a 4 by 4 linear system of can be written:

0

Dissolved Oxygen (μMol/kg)

fAtl ½S Atl þ fPac ½S Pac þ fMet ½S Met þ fSI ½S SI ¼ ½S Obs fAtl d18 O Atl þ fPac d18 O Pac þ fMet d18 O Met þfSI d18 O SI ¼ d18 O Obs

Depth

AOS-2005: All Samples

(2)

fAtl ½N Atl þ fPac ½N Pac þ fMet ½N Met þ fSI ½N SI ¼ ½N Obs fAtl þ fPac þ fMet þ fSI ¼ 1

400 100 350

200

300 300 400 250 0

where S = salinity, and [N] = nutrient concentration, and the last equation is the conservation of mass [see, e.g., Ekwurzel et al., 2001]. The system can be solved in matrix form: 1

½ f ¼ fC g ½ y

500

15

Pot. Dens.

AOS 2005: All Samples

b

1028

(3) 15

1027

Nitrate (μmol/kg)

⋯ ⋱ ⋯

39 SSI = ⋮5 ; 1

10

Nitrate (μMol/kg)

where f is the vector of component fractions, y is the vector of observations for a sample, and C, the design matrix, is composed of the concentrations in the end members from equation 2: 82 < SAtl fC g ¼ 4 ⋮ : 1

5

(4)

1026 10 1025 1024 5 1023 1022

5.2.2. Pacific Water Tracer [32] To approximately account for biological processes, water mass studies have used the composite tracers: dissolved oxygen 1:95 175

N ¼ 0:87 ðnitrate ð16 phosphateÞ þ 2:9Þ NO ¼ nitrate þ

dissolved oxygen 11

0.5

distance to Atlantic fit ðdistance to Pacific fitÞ þ ðdistance to Atlantic fitÞ

(6)

[34] ANP is essentially a version of the N* tracer, adapted to the specific N/P ratios of the Arctic water column and scaled to the dynamic range of the pelagic Arctic Ocean. In the past, we have used the globally defined PO* tracer, which is based on an average Redfield ratio of 175 : 1 for oxygen to phosphate [Broecker et al., 1985, 1995]. However, observations indicate that above the Atlantic layer in the Arctic, the observed O/P ratio is significantly lower than the global average of 175, while below the halocline, in water that is mostly advected from the North Atlantic, the global averages are a good match for regenerated nutrients.

2

20

(7)

(8)

1.5

Red: Chukchi Sea, Pink: Canada Basin Green: Bering Sea ( 0.99; p < 1e4), and where they disagree, we trust the local definition, ANP, over the globally defined N*. We select the POs* tracer over NO because the system is nitrate limited, whereas there is significant phosphate in all of our samples, so that phosphate is sensitive to water mass composition gradients in some locations where nitrate is depleted. [36] Before proceeding to the results of the water mass decomposition, we would like to comment on the ANP “Pacific” end-member. Figure 3C shows in the N : P space for data from several cruises. The Bering Strait inflow is intermediate between the Atlantic and Pacific end members [Codispoti et al., 2005; Yamamoto-Kawai et al., 2006] (Figure 3c); i.e., the ANP Pacific end-member is Pacific inflow modified by the addition of phosphate and/or the loss of nitrate as the inflow transits the broad continental shelf in the Chukchi Sea. [37] East Siberian Sea water (ESW), with high runoff concentrations from the Lena and smaller Siberian rivers, has nitrate and phosphate concentrations of approximately 0.6 and 1.4 mM, respectively [Nitishinsky et al., 2007], and flows toward the study region [Weingartner et al., 1999]. With its low N/P ratio, a mixture of 25% ESW and 75% Bering Strait inflow (BSI) would fall along the ANP end-member in N-P space. On the other hand, Devol et al. [1997] report that the denitrification rate in sedimentary pore waters of the Chukchi Sea is approximately 1 mg atom m2 day1 (1 mole weight of nitrogen atoms per square meter per day). Equivalent to approximately 5 mM nitrate drawdown per square meter per year, this process could also contribute to the transformation, depending on the residence time of BSI on the continental shelf. [38] In a previous study of water masses over the pelagic Arctic Ocean, Ekwurzel et al. [2001] identified the waters detraining from the Chukchi Shelf into the Canada Basin as Mixed East Siberian and Chukchi Shelf water (MECS). We follow their nomenclature below. However, we note that the largest portion of this water must come from the Pacific

Ocean, which is both modified through contact with the bottom sediments and atmosphere, and diluted by mixing. [39] To ascertain the most likely mixing ratio of BSI and ESW, we tested combinations of BSI and ESW against a cost function that takes into account the nutrient concentrations, salinity, and residence time on the Chukchi shelf. We tested mixing ratios in 1 increments between 75% and 100% BSI. For each case, we assumed a uniform denitrification rate of 1mg atom m2 day1. From mass balance and tracer age considerations, we restricted the mean residence time of BSI on the Chukchi shelf to between 9 and 18 months. Our optimization minimized the number of samples showing non-physical water mass fractions (1) after decomposition with the ANP end-member method. The optimal mixing ratio by these criteria was 85% Pacific inflow and 15% “indigenous” Arctic shelf water. [40] We note that given the uncertainty of the water-mass inversion technique, some fractions above one or below zero are expected. For example, of 536 samples shallower than 500 m, there were 92 samples with MECS fraction below zero (average fraction 0.007) and 42 samples with MECS fraction above 1 (average fraction 1.0151). Assuming that there are some samples in which MECS, Atlantic, or meteoric water is essentially missing, these high and low values are well within the expected errors of the water-mass inversion technique. [41] The POs* pattern (Figure 2i) is similar to that of phosphate, except that it fills in the region above the nutrient maximum tongue with higher values and more horizontal structure. In water mass terms, this indicates that the region above the nutrient maximum tongue also contains a large fraction of MECS and that the low-nutrient values toward the surface are, as expected, a result of photosynthetic processing. [42] The ANP pattern (Figure 2H) is structurally very similar to POs*. The main difference is that the ANP tracer is more bimodal, with samples clustering near either zero or 1. By contrast, more of the POs* values are intermediate between the POs* minimum and maximum (approximately 0.6 and 2.17 mmol kg1. [43] Both ANP and POs* are impacted by processes other than photosynthesis and respiration, which cause departures from Redfield ratios and account for divergence between the ANP and POs* composite tracers. In the case of ANP, the principal non-Redfield factors are bacterial nitrification and denitrification, which take place mainly in the anoxic regions of the continental shelf benthos. For POs*, the main divergences from Redfield ratios are caused by the addition of phosphate from sediments, carried either in river runoff or pore waters, and the equilibration of dissolved oxygen at the surface. Since the processes impacting ANP and POs* are independent, we take the similarity of their distributions as validation that the divergence from local Redfield ratios is minor. While acknowledging the errors inherent in using biological processing as water mass tracers, we note that they remain the best tools available to us for distinguishing Pacific from Atlantic waters. [44] Jones et al. [1998] used the AOS94 data to define the ANP method, and by definition, the AOS94 samples along the “Pacific” N/P relationship were considered to have no Atlantic water. Analyzing the 1994 data using the ANP method of Jones et al. yields peak Pacific water

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fractions over 90%. Ekwurzel et al. analyzed these same samples using the PO* method and got peak Pacific water fractions below 60%. The discrepancy in peak values is largely eliminated in this reanalysis of the 1994 samples by (1) identifying arctic-specific ratios of oxygen to phosphate in waters with salinities below 34.5, (2) re-assessing the end-member concentrations, and (3) quantifying the mixture of Bering Strait Inflow and indigenous shelf waters that comprises MECS. 5.2.3. End-Member Choices [45] The choice of end members has been discussed in detailed elsewhere [see, e.g., Newton et al., 2008; Ekwurzel et al., 2001, and references therein]. For salinity, we use the following: Atlantic Water: 34.92, Pacific Water: 32.5, Sea-ice meltwater: 4 1, and river runoff: 0. [46] For oxygen isotope ratios (H2 18O/H2 16O), we use the following: Atlantic Water: 0.3, river runoff: 18 2, and Pacific water: 1.1 0.2 per mil. We have followed the practice of Bauch and co-workers [1995, 2012] and Ekwurzel et al. [2001] in assigning a sea-ice d18O value that is 2.6 per mil higher than that of the water at the surface of each station. This strategy is based, in part, on field observations by Melling and Moore [1995]. Macdonald et al. [1999] used contemporaneous sea-ice and surface water measurements from the SHEBA Drift stations (1997–1998) to evaluate this strategy and also found it appropriate. Similarly, comparison of sea-ice motion to d18O profiles confirms that the d18O of basal sea ice is consistently offset from that of the surface water over which it has travelled [Pfirman et al., 1997]. [47] For POs*, we use the following: Atlantic Water: 0.7 0.05, shelf-modified Pacific water: 2.4 0.26, meteoric water: 0.1 0.1, and sea ice is set to the value in the surface waters at each station. The meteoric water values are taken from Ekwurzel et al. [2001]. Varying these settings has very little impact on the water mass calculations, as meteoric and sea-ice fractions are essentially set by the salinity and d18O values, and their contributions to the POs* balance are minimal. [48] Ekwurzel et al. [2001], Schlosser et al. [2002], and Chen et al. [2008] assigned a Pacific end-member value of POs* of 2.4 0.3, which we adopt here as well. We note that the success of this value in the current study is somewhat fortuitous, since we use an O : P ratio of 125 for water with salinity less than 34.5. Thus, our choice of POs* = 2.4 implies that the Pacific end-member has less phosphate than the end-member chosen for earlier studies. In fact, Ekwurzel et al. [2001] based their choice on a small number of nutrient measurements reported in Codispoti and Richards [1968] and Bjork [1990] in the northern Bering Sea, at the “headwaters” of the inflow through Bering Strait [Ekwurzel et al., 2001]. Since that time, there have been many high-quality measurements in the Bering Sea as part of the BEST program and in the Chukchi Sea as part of the SBI program. Both data sets exhibit lower O : P ratios than the global average for salinities below about 34.5, supporting the ratio of 125 : 1 used here. Using this lower ratio, the BEST data north of 60 N and over a depth horizon of 40 to 300 m (175 samples) have a POs* value of 2.62 0.13. The SBI data within a salinity band of 33 and 33.2 (170 samples), i.e., waters that are on the isopycnal of the peak of the UHW nutrient maximum, have a POs* value of 2.36 0.27. This

decrease in POs* may be linked to oxygen equilibration at the surface, but we note that it is also consistent with a Bering Strait inflow diluted with 13% of Atlantic water, similar to the 15% dilution we found for the shelf-modified Pacific inflow end-member when optimizing for the ANP end-member. We use the slightly larger value 2.4 because while consistent within errors with the SBI data, it minimizes the number of non-physical fractions in the deep layers along the AOS-05 cruise track. This end-member, based on the SBI data, also represents shelf-modified Bering Strait Inflow, and the POs*-based results should be directly comparable to those based on the ANP tracer (Table 1). 5.2.4. Comparison With Alkalinity-based Results [49] Jones et al. [2008] performed a similar decomposition of samples from the AOS 2005 cruise, using alkalinity in the place of oxygen isotope ratios to distinguish river runoff from sea-ice meltwater. An advantage of alkalinity is that it can be analyzed on board a research vessel and is available at the end of a cruise. A disadvantage is its wide variability in natural waters, as a result of its link to ion concentrations, which are impacted by biological processing. [50] There are several other differences in our approaches. We use a reference salinity of 34.92, whereas Jones et al. used 34.85. Jones et al. assigned the Atlantic N/P relationship to sea-ice meltwater, whereas we have used the surface water value for nutrient tracers at each station. They assume a sea ice salinity of 2.5, whereas we assume 4 psu. Thus, there is no reason to expect quantitative agreement between our results and theirs. However, in qualitative terms, the two methods generally agree. [51] Our Pacific Freshwater distribution is nearly identical to Jones et al. [2008] when we use the same nutrient tracer and slightly different, as discussed above, when we use POs*. In qualitative terms, their River Fraction and our Meteoric Water Fraction fit closely. Where there is significant runoff content, our Meteoric Water values tend to be between 10 and 20% higher than theirs for River Runoff. Peak values in the runoff plumes over the southern Canadian Basin and along the Lomonosov Ridge are approximately 16% in our analysis and 13% in theirs. Our sea-ice meltwater values also show the same features as those found by Jones et al. [2008]. However, our values are consistently lower, between 8% (8% brine formation) and +2%, versus 4% to +6% in the Jones et al. analysis. 5.3. Freshwater Component Distributions [52] The above equations and end members were used to calculate the freshwater components in the upper 500 m along the AOS05 section. The individual components add up to the total freshwater anomaly along the AOS 2005 track (Figures 2c and 5a), relative to the salinity of the Atlantic

Table 1. End-Member Parameter Values

Atlantic Pacific Meteoric Sea-ice Melt

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Salinity (PSU)

d18O (Per Mil)

POs* (mmol kg1)

ANP (Fraction)

34.92 32.5 0 4

0.3 1.1 20.0 Surface + 2.6

0.7 2.4 0.1 Surface

0 1 0 Surface


200 m depth in the Canada Basin; and the subsurface mesoscale eddy at Station 4 over the Alaskan continental slope. [53] The pooling of freshwater in the Canadian Basin and its relative absence from the Eurasian side of the LR are a consequence of a generally positive vorticity input over the Eurasian Basin and negative input over the Canadian Basin. Accumulating surface (fresh) water over the BG causes downward Ekman pumping and downward bowing of the freshwater isolines there. Corresponding upward Ekman

end-member (34.92). In the Arctic, the freshwater content is a close proxy for the specific volume anomaly, a, so features of the geostrophic circulation oriented perpendicular to the transect are also visible: Atlantic inflow along the Eurasian (right hand) boundary; a plume of relatively fresh water moving from Siberia to Fram Strait (out of the page in Figures 2 and 4) on the Eurasian flank of the LR; the wide Beaufort Gyre over the Canadian Basin, with isopycnals descending from the surface of the Makarov Basin to about

e 20 10 0

Height (m3/m 2)

Depth (db) Depth (db)

Sea-ice Meltwater (%, Negative = Brine)

20

Height (m3/m 2)

Integrated Meteoric Water Fraction

b

c

Integrated Freshwater Anomaly: Upper 500 m Height (m3/m 2)

Freshwater (% Anomaly from S = 34.92) 4 8 12 20 30 50 52

Depth (db)

a

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Major Ridge Systems Major Basins (AOS05 Track) 1000 Canada Canada Makarov Nansen 2000 3000 4000 0

i

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Lomonosov Gakkel

3000 4000 0

500 1000 1500 2000 2500 Distance (km) from Barrow, Alaska

500 1000 1500 2000 2500 Distance (km) from Barrow, Alaska

MECS from the Arctic N/P method (%)

Figure 4. Water-mass fractions calculated using the PO4* composite tracer of Pacific inflow. (left) Water masses along the AOS 2005 transect. (White space in upper right is missing data.) (right) Freshwater components vertically integrated over the upper 500 meters. Sea-ice freshwater fraction = 0.8855 * sea-ice fraction. Modified East Siberian Chukshi Shelf (MECS) freshwater fraction = 0.0693 * MECS fraction. Note that scales are different for each chart on the right. (bottom-left) MECS from the Arctic N/P composite tracer. 2142


pumping on the Eurasian side keeps relatively saline water at the surface there [Hunkins and Whitehead, 1992; Newton et al., 2008; Proshutinsky et al., 2002, 2009]. [54] The geostrophic flow in the upper 500 m driven by the freshwater distribution would be from Siberia toward Greenland (out of the page in Figures 2 and 5) in a very broad region from the Beaufort Gyre in the Canadian Basin approximately to the Lomonosov Ridge. This pattern is reflected in the average sea-ice tracks as well, e.g., [Zhao and Liu, 2007]. 5.3.1. Meteoric Water [55] The largest meteoric water concentrations are observed in the center of the BG and along both sides of the Lomonosov Ridge, where the meteoric water concentrations peak at about 17% of the water mass (Figure 4b). (River Runoff distributions using the POs* and ANP methods are nearly identical (r = 0.999, even after excluding the waters below 300 m), so only the POs*-based result is shown). Meteoric water fractions decay smoothly with depth, to about 12 % at 500 m. Although the concentrations are low, the deep waters below the Atlantic Layer hold approximately half the full water-column burden of freshwater. It is this addition of freshwater mainly made up of river runoff and MECS that accounts for the salinity difference between Canadian Basin Deep Water (CBDW) and the waters of the North Atlantic. The major fraction of the meteoric water in the BG has been shown elsewhere, using barium concentrations, to be of North American origin, mainly from the Mackenzie and Yukon rivers (the latter via the Bering Strait, e.g., Guay and Falkner [1997]). Modeling results [e.g., Newton et al., 2008] also indicate a significant fraction of Eurasian river runoff entrained in the gyre, depending on large-scale wind patterns. The meteoric water peak at station 25, over the Makarov Basin and the peak over the LR are more likely from Eurasian rivers, the Lena, Ob, and Yenisei being the largest. The runoff from those rivers travels eastward, along the Siberian shelf seas, and detrains from the continental shelves preferentially along the topographic saddles where the major ridges meet the shelf [Newton et al., 2008; Anderson et al., 2004; Weingartner et al., 1999; Bauch et al., 1995, 2009, 2011]. 5.3.2. Sea Ice Meltwater [56] Sea-ice meltwater may be either positive or negative, with the “negative” sea-ice melt indicating the fraction of brines from net sea-ice formation over the history of a water

Arctic N/P−Based Analysis

0.15

b 1 0.8 0.6

0.1

0.4 0.05 0.2 0 0

0.05

0.1

0.15

0

Arctic N/P−Based Analysis

Meteoric Water Fraction POs * vs ANP Methods

a

parcel. Since sea ice is exported from the Arctic through Fram Strait and (to a lesser extent) the Canadian Arctic Archipelago, Arctic waters must, on average, exhibit net sea-ice formation. Along the AOS-05 cruise track, there is a very thin layer of net positive sea ice melt at the surface over the Alaskan continental shelf, over parts of the BG, and at the final station, 53, over the Barents Sea continental margin (Figure 4c). The Nansen Basin, especially directly north of the Barents Sea, is known from previous studies to have low runoff fractions and positive sea-ice meltwater content down to a depth of about 150 m [Bauch et al., 1995; Ekwurzel et al., 2001; Schlosser et al., 2002; Newton et al., 2008]. Thus, the sea-ice meltwater at the top of Station 53 appears to be a persistent feature. The sea-ice meltwater layer at the surface of the Canadian Basin, on the other hand, may be seasonal. By late August, there would be a thin (~10 m) summer mixed layer which would be heavily influenced by that summer’s basal melting. In the winter, a deeper winter mixed layer forms, which is similarly impacted by brines rejected during sea ice formation. [57] Between the Alaskan shelf and the AMR, maximum sea-ice formation coincides with the nutrient-rich peak in shelf-modified Pacific inflow. This reflects the fact that sea-ice formation is linked to the transformation of the Pacific inflow (at an average salinity of about 32.5) into the water mass that detrains from the Chukchi shells at the base of the upper halocline waters (UHW). The brine and nutrient peaks are centered about a salinity of approximately 33.1. We have seven samples from the maximum sea ice formation plume (salinity between 33 and 33.2; average salinity = 33.14), which exhibit the nutrient characteristics of MECS (averages: phosphate = 1.87 mmol kg1; nitrate = 15.42 mmol kg1; POs* = 2.17 mmol kg1). Chukchi Sea waters exhibit a great deal of scatter in nitrate-phosphate-salinity space and in principle must exit the shelf with a broad range of densities and nutrient signatures. However, the spatially consistent peak centered at a salinity of 33.1 indicates that most of these small-scale gradients are mixed away during and following the detrainment process. From the peak, MECS is mixed upward, into the upper halocline as well as downward into the lower halocline between the MECS peak and the Atlantic layer. [58] In addition to salt from brine formation, the peak of the MECS tongue contains a significant contribution of salinity from mixing with approximately 11% Atlantic

0.1

Sea−ice Melt Fraction POs * vs ANP Methods

0.05 0 −0.05

−0.1

0

0.1

POs*−Based Analysis

POs*−Based Analysis

Figure 5. Comparison of phosphate-star (horizontal axis) and Arctic N-P-relationship (vertical axis) methods for estimating meteoric (left) and sea-ice melt (right) fractions. 2143


water. The profiles at Station 4, where we believe there is an anti-cyclonic mesoscale eddy, suggest one potential pathway for mixing Atlantic water into the halocline. The 33.1 salinity horizon is deeper, at 212 m, and has about 15% Atlantic water content. Here the 33.1 isohaline sits in the upper part of an abrupt transition zone between the main body of the eddy, which extends down to 200 m, and the Atlantic Layer, at 260 m. It is possible that the eddy is effectively “pumping” Atlantic water up into the 33.1-salinity isopycnal, via shear-induced mixing at the edges of the eddy. As the eddy winds down, some Atlantic water would detrain along the 33.1 isohaline. 5.3.3. Mixed East Siberian and Chukchi Shelf Water (MECS) [59] Qualitatively, the identification of Pacific inflow is clear (Figure 4d). A tongue of high-nutrient waters extends across the upper waters of the Canadian Basin, which identifies it as having a high concentration of North Pacific water. Shelf-modified Pacific water is a significant fraction of the water in the upper halocline and surface waters throughout the Beaufort Gyre, down to a depth of about 300 m at the center of the gyre, and following isohalines up and to the North until it trails off over the Makarov Basin. In addition, there is a narrow plume of water on the Eurasian side of the Lomonosov Ridge that has a similar nutrient signal. This plume is transported by the Trans-Polar Drift (TPD), which flows north along the Ridge from Siberia to Greenland. The high POs values in the plume, and the high ANP values, identify it as having a high pre-formed phosphate content and also being highly denitrified, relative to the Atlantic water N/P ratio. However, given the circulation pattern along the LR, we doubt that this chemical signal originates with Bering Strait Inflow that has been modified over the Chukchi or East Siberian shelf. Rather, we think this is more likely a plume of mixed Laptev Sea and East Siberian Sea shelf water. Bauch et al. [2011] come to the same conclusion in analyzing an extensive data set taken in 2007 over the Eurasian sector. [60] Quantifying the Pacific fraction is made somewhat ambiguous by the difficulties identified earlier with respect to biological processing of nutrients and the choice of end-member. We present the Pacific inflow, which we term Mixed East-Siberian Chukchi Sea water (MECS), content for two nutrient composites: POs* (Figure 4d) and ANP (Figure 4i). In general, the POs* method yields lower estimates of the MECS fraction than the ANP method, and the POs* method has more structure within the MECS-maximum zone. The meteoric water fractions derived from the two methods diverge as a function of the volume of MECS in the sample (Figure 5a). In the regions with significant MECS contributions to the overall water mass composition, the POs* method estimates a higher fraction of Pacific water than the ANP method. Since Pacific inflow has a lower salinity than Atlantic water, the POs* calculation compensates with a lower meteoric water input than the ANP method. Sea-ice meltwater fractions are nearly identical for both methods (Figure 5b). 5.4. Freshwater Component Water Column Inventories [61] A convenient way to visualize the horizontal distribution of freshwater components is to vertically sum their concentrations to get the water column content.

Conceptually, this is equivalent to taking all of the water of a given type out of the water column and stacking it at the surface. Figure 4e shows the freshwater inventory of the upper 500 meters along the Oden 2005 track, while in Figures 4f–4h, we depict the decomposition of the total freshwater inventory into fractions of meteoric water, seaice meltwater, and Pacific Water, respectively. With respect to the reference water mass, Atlantic inflow with a salinity of 34.92, freshwater content is calculated as (S0 S)/S0, 6.93% and 88.55% for MECS and sea-ice meltwater, respectively. Thus, the “Pacific freshwater” content of each sample is 0.0693 times its MECS content and likewise for the sea-ice freshwater. In addition to providing a summary visualization of the horizontal distribution of freshwater components, the vertical sums are integral quantities useful for comparisons across regions or timeframes. 5.5. AOS05 Station 4: A Slope-Trapped Eddy? [62] Station 4 samples an anomalous feature over the continental slope. At depths between about 50 and 250 m, salinity and consequently potential density are lower than at either neighboring station (Figure 6a). At about 50 m, the density (and salinity) profiles cross so that close to the surface the densities at Station 4 are higher than at the neighboring stations. Relative to the 500 m isobar, the picture is of a baroclinic eddy, rotating anticyclonically between the surface and about 100 m, and cyclonically below that with the cyclostrophic velocities gradually diminishing with depth and disappearing somewhere between 300 and 350 m. [63] The core of the feature is warmer than neighboring stations, by as much as 8 C in the upper 160 m (Figure 6b). Between about 180 and 350 m, it is slightly cooler than nearby stations. The nutrient peak, endemic to MECS water and characteristic of the Upper Halocline Layer, is depressed, relative to surrounding stations, by about 50 m, from 150 to 200 m (Figures 6c and 6d). Water-mass decomposition confirms the increased depth of the MECS peak and also indicates that between 50 and 160 m, there is no net seaice formation, while at Station 5, the water contains about 2.5 to 3 % of brine from ice formation (Figure 6f). [64] Nitrate and phosphate are both low in the upper 100 m, and peak at 200 m. (Figures 6c and 6d), about 80 m below the nutrient peaks at Stations 3 and 5. Between the surface and the nutrient peak, there is a deficit in nitrate, whereas below the peak, the phosphate declines, leaving a nitrate excess. Nutrient distributions are similar to the surrounding waters when viewed as a function of salinity. Also, phosphate and nitrate from Station 4 overlay closely on the balance of the cruise data in nitrate-phosphate-salinity space. However, the isohalines within the eddy are depressed by about 80 m, relative to ambient Canada Basin waters. The overall impression is of a dynamical feature that has suppressed the halocline by about 80 m and induced a certain amount of vertical mixing in the process. [65] The largest mismatch between the “ANP” and “POs*” methods of water-mass estimation occurs in the upper portion of this eddy. The eddy does not stand out in the ANP-based analysis, because the nitrate and phosphate concentrations co-vary there. POs*, on the other hand, is influenced by the relatively low dissolved oxygen concentrations in the upper part of the eddy (Figure 2j). POs* is sensitive to oxygen levels and yields a lower estimate of MECS

2144


a

b

100

Depth (db)

Depth (db)

100

300

300

1020

1022

1024

1026

1028

−2

0

Potential Density (kg/m3)

2

4

6

8

Temperature (oC)

c

d

100

Depth (db)

Depth (db)

100

300

300

0

5

10

15

0.5

1

Nitrate (μMol/kg)

1.5

2

Phosphate (μMol/kg)

f

Freshwater fractions: Stations 4 (solid), 5 (open) Magenta: sea ice, Green: meteoric, Blue: MECS

e

100

Depth (db)

Depth (db)

100

300

300

−4

−3

−2

−1

0

1

0

δ18O (per mil)

0.2

0.4

0.6

0.8

1

Water Components (fractions)

Figure 6. Water properties and water-mass decomposition at AOS 2005 stations 3 (green) 4 (red) and 5 (blue). In Figure 6f, Station 4 samples are solid dots and Station 5 are unfilled squares.

than the ANP method in the upper portion of the eddy. Oxygen is not under-saturated in the eddy (Figure 2e), indicating that the low dissolved oxygen values are likely a result of the high temperature of its source waters. This demonstrates a limitation of the POs* tracer, which is that it is not conservative at the surface, where there is gas exchange with the overlying atmosphere. [66] The feature at Station 4 is quite different from the cold-core eddy that was surveyed in detail in roughly the same area in 1997 [Muench et al., 2000]. That eddy exhibited an excess of brine and was denser than its surroundings. Muench et al. concluded that it was likely the product of a sea-ice formation event, probably at a near-shore polynya on the Chukchi shelf. The Station 4 feature could not have been

created by brine-driven convection; its core is less saline and therefore less dense than its surroundings. The water-mass inversion calculation indicates less sea-ice formation at Station 4 between about 50 and 180 meters, where the oxygen isotope ratios are vertically homogeneous (Figure 6f). [67] The heat content of the core identifies it as likely having formed over the continental shelf during summer. Like the 1997 eddy, this feature has nutrient relationships indicative of MECS, but the core of the nutrient peak lies below the warm core visible in Figure 6b. [68] The feature must have been formed in summer, most likely over the East Siberian or Chukchi Shelf in waters that had been freshened by sea-ice melt. It probably extended from the surface to the bottom and was forced northward

2145


across the continental shelf break by winds and/or ambient currents. Deepening of the feature would have created a more compact, more intense cyclostrophic feature through vortex stretching. Once over the continental shelf, the motion of least resistance would be along isobaths and from West to East. As it migrated along the shelf, it would be subject to mixing as it slowly wound down. Vertical mixing is evident in the profiles between roughly 200 and 300 m, probably driven by vertical shear instabilities, and would

have entrained a fraction of Atlantic Water. The most rapid mixing would have been along isopycnals, which slope upward from Station 4 to both the North and the South. [69] If this speculation is correct, then the feature sampled at Station 4 has travelled along a similar pathway as the cold-core eddy observed in 1997 but began its trip at the peak of summer, as opposed to the fall freeze-up. Its vertical extent is similar to the 1997 eddy. Our horizontal resolution is insufficient to assess its width but assuming it is similar, it

d Modified E. Siber/Chukchi Sea water 0

Depth (db)

100 200 300 400 0

0.04

0.06

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Depth (db)

Depth (db)

a Salinity

0.02

200 300

200 300 400

400 31

32

33

34

35

Salinity (psu)

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0.02

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0.1

f Sea-ice Melt (Formation) freshwater 0

b Potential temperature 0

100

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g Makarov Basin Freshwater Anomaly 2005 average − 1994 average 0

c Freshwater 0

100

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100 200 300

Sea−ice melt: Red Meteoric water: Blue Sum: Black

200 300 400

400 0

0.02

0.04

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Freshwater content (S − S0)/S0

−0.01

0

0.01

0.02

Fraction

Figure 7. Comparison of properties and water mass estimates over the Makarov Basin: Black: AOS2005, Stations 26 through 33, 36 and 39. Red: AOS-1994, Stations 26 through 30 (Station locations marked with circles in Figure 1.). 2146


seems likely that the eddy transport mechanism discussed by Muench et al. [2000], holds for summer, as well as winter, shelf waters. 5.6. Freshwater Inventories Over the Markov Basin: 1994 versus 2005 [70] The 2005 and 1994 Arctic Ocean sections both cross the Makarov Basin, albeit at somewhat different locations. In 2005, 10 tracer profiles were taken in the Makarov Basin, stations 26 through 33, 36, and 39; in 1994, there were five profiles at comparable locations, stations 26 through 30. We do not include several Makarov stations that lie over the continental slope, as those are in the Atlantic boundary current, and have a different hydrographic structure. Figure 7a shows the salinity at all 15 stations, and the average of the 1994 and 2005 profiles. Figure 7b shows the analogous plot for the potential temperature (see also Swift et al. [1997] and Ekwurzel et al. [2001] for detailed discussion of the 1994 data.) [71] The surface temperature in both years was pinned at the freezing point. There is a thin Cold Halocline layer, between about 25 and 55 m, in which salinity and density are highly stratified, but temperature remains within a few tenths of a degree of the freezing point. Below the Cold Halocline, the temperature increases rapidly to the Atlantic Layer maximum between about 250 and 325 m. Below the Atlantic Layer (not shown), the potential temperatures converge and are not statistically different within the depth range below about 1000 m. However, in the Atlantic layer, the 2005 data indicate that the water is warmer by up to 0.4 C than in 1994. The difference in heat content, integrated through the water column, amounts to approximately 460 MJ m2. The heat difference between 2005 and 1994 is sufficient to melt a layer of about 1.4 m of sea ice, essentially the entire summer sea ice cover over most of the Makarov Basin. [72] Salinity values scatter widely in the mixed layer. Below the depth of the winter mixed layer, on the other hand, the profiles show little variation, either spatially or between the two cruises. The breakdown of the freshwater anomaly into its components (Figures 7c–7f) indicates that in 2005, the surface waters of the Makarov Basin contain more sea-ice meltwater, less meteoric water, and less Pacific inflow than in 1994. (Note that the average of the total freshwater inventories is not an exact match to the sums of the analogous freshwater component inventories. For each bottle, the sums of freshwater components is an exact match to the total freshwater, which is calculated directly from the salinity. However, there are many more salinity measurements from the two cruises than there are d18O measurements, including at several stations where oxygen isotopes were not measured at all. This difference in resolution introduces about an error of approximately 1.5% in the overall basin inventories). Errors on the MECS fraction, largely as a result of non-conservative behavior of nutrients and gas exchange at the surface, are large enough that the difference between 1994 and 2005 is not significant. The standard error on each estimate of meteoric and sea-ice meltwater fractions, due to uncertainty in the end-member estimates and analytical errors, is about 2%. Vertical interpolation adds about 1.5% to the standard error, while averaging over several profiles drives the errors down by about a factor of 0.4, resulting in a standard error of approximately 2% on the averaged profiles presented in Figures 7e and 7f. Figure 7g shows

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Figure 8. Seasonal ice zone from satellite imagery. the differences in freshwater (black), meteoric water (blue), and sea-ice melt fractions between 1994 and 2005. The differences in meteoric water and sea-ice melt at individual depths are not statistically significant, but the inventories, representing several dozen measurements, are. 5.6.1. Sea-Ice Meltwater Concentrations [73] The warming of the Arctic has led to an increase in the amplitude of the annual sea-ice melt/freeze cycle, which is reflected in the expansion of the Seasonal Ice Zone (Figures 8a and 8B). More melt in summer increases the concentration of meltwater at the surface. More sea ice formation in winter increases brine-driven convection and

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the concentration of brines at the base of the winter mixed layer. This pattern of sea-ice meltwater differences is reflected in Figure 7g. Sea-ice melt (the red line) has a minimum at the base of the winter mixed layer, and a peak at the surface is consistent with such a shift in dynamics. As discussed above, the size of the differences,