A Tale of Two Basins an Integrated Physical and Biological Perspective of the Deep Arctic Ocean 2015...

Progress in Oceanography 139 (2015) 89–121

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Progress in Oceanography

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A tale of two basins: An integrated physical and biological perspectiveof the deep Arctic Ocean

http://dx.doi.org/10.1016/j.pocean.2015.07.0110079-6611/� 2015 Elsevier Ltd. All rights reserved.

⇑ Corresponding author at: UiT - The Arctic University of Tromsø, Department of Arctic and Marine Biology, 9037 Tromsø, Norway. Tel.: +47 776 44382.E-mail address: [email protected] (B.A. Bluhm).

B.A. Bluhm a,b,⇑, K.N. Kosobokova c, E.C. Carmack d

a UiT - The Arctic University of Tromsø, Department of Arctic and Marine Biology, 9037 Tromsø, Norwayb School of Fisheries and Ocean Sciences, Institute of Marine Science, University of Alaska Fairbanks, 905 Koyukuk Drive, Fairbanks, AK 99775-7220, USAc P.P. Shirshov Institute of Oceanology, Russian Academy of Sciences, Nakhimovsky Prospect 36, Moscow 117218, Russiad Fisheries and Oceans Canada, Institute of Ocean Sciences, 9860 West Saanich Road, Sidney, British Columbia V8L 4B2, Canada

a r t i c l e i n f o

Article history:Available online 8 August 2015

a b s t r a c t

This review paper integrates the current knowledge, based on available literature, on the physical andbiological conditions of the Amerasian and Eurasian basins (AB, EB) of the deep Arctic Ocean (AO) in acomparative fashion. The present day (Holocene) AO is a mediterranean sea that is roughly half continen-tal shelf and half basin and ridge complex. Even more recently it is roughly two thirds seasonally and onethird perennially ice-covered, thus now exposing a portion of basin waters to sunlight and wind. Basinboundaries and submarine ridges steer circulation pathways in overlying waters and limit free exchangein deeper waters. The AO is made integral to the global ocean by the Northern Hemisphere ThermohalineCirculation (NHTC) which drives Pacific-origin water (PW) through Bering Strait into the Canada Basin,and counter-flowing Atlantic-origin water (AW) through Fram Strait and across the Barents Sea intothe Nansen Basin. As a framework for biogeography within the AO, four basic, large-scale circulation sys-tems (with L > 1000 km) are noted; these are: (1) the large scale wind-driven circulation which forces thecyclonic Trans-Polar Drift from Siberia to the Fram Strait and the anticyclonic Beaufort Gyre in the south-ern Canada Basin; (2) the circulation of waters that comprise the halocline complex, composed largely ofwaters of Pacific and Atlantic origin that are modified during passage over the Bering/Chukchi andBarents/Siberian shelves, respectively; (3) the topographically-trapped Arctic Circumpolar BoundaryCurrent (ACBC) which carries AW cyclonically around the boundaries of the entire suite of basins, and(4) the very slow exchange of Arctic Ocean Deep Waters. Within the basin domain two basic water massassemblies are observed, the difference between them being the absence or presence of PW sandwichedbetween Arctic Surface Waters (ASW) above and the AW complex below; the boundary between thesedomains is the Atlantic/Pacific halocline front. Both domains have vertical stratification that constrainsthe transfer of nutrients to the surface layer (euphotic zone), thus leading to their oligotrophic state, par-ticularly in the more strongly stratified Pacific Arctic where, despite high nutrient values in the inflow,convective reset of surface layer nutrients by haline convection in winter is virtually absent. First andmulti-year sea ice drastically alters albedo and insulates the underlying water column from extreme win-ter heat loss while its mechanical properties (thickness, concentration, roughness, etc.) greatly affect theefficiency of momentum transfer from the wind to the underlying water. Biologically, sea ice algal growthin the basins is proportionally almost equal to or exceeding phytoplankton production, and is a habitatand transport platform for sympagic (ice-associated) fauna. Owing to nutrient limitation due to strongstratification and light limitation due to snow and ice cover and extreme sun angle, primary productionin the two basin domains is very low compared to the adjacent shelves. Severe nutrient limitation andcomplete euphotic zone drawdown in the AB favors small phytoplankton, a ubiquitous deep chlorophyllmaximum layer, a low f-ratio of new to recycled carbon fixation, and a low energy food web. In contrast,nutrients persist –albeit in low levels– in the western EB, even in summer, suggesting light limitation,heavy grazing or both. The higher stocks of nutrients in the EB are more conducive to marginal ice bloomsthan in the AB. The large-scale ocean currents (NHTC and ACBC) import substantial expatriate, not locallyreproducing zooplankton biomass especially from the adjoining subarctic Atlantic (primarily Calanus fin-marchicus), but also from the Pacific (e.g., Pseudocalanus spp., Neocalanus spp. and Metridia pacifica). Theseadvective inputs serve both as source of food to resident pelagic and benthic biota within the basins, and

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Nomenclature

AcronymsAB Amerasian BasinAO Arctic OceanAOI Arctic Oscillation IndexACBC Arctic Circumpolar Boundary CurrentASW Arctic Surface Waters

AW Atlantic-origin WaterBSB Barents Sea BranchEB Eurasian BasinFSB Fram Strait BranchNCP net community productionPW Pacific-origin water

90 B.A. Bluhm et al. / Progress in Oceanography 139 (2015) 89–121

as potential grazers exerting top down control on limited phytoplankton resources. Benthic organismswithin the AO basin show previously unappreciated biodiversity and surprising dispersion of speciesgiven the isolation of individual basins and low vertical carbon flux and resulting biomass. Larval disper-sion is aided by the large-scale flows and perhaps, we hypothesize in the deep benthos by convectiveupdrafts driven by geothermal heating. Zooplankton diversity, in contrast, is low, but again faunal assem-blages are equally distributed between the EB and AB. Species pools of both pelagic and benthic commu-nities change more with water depth rather than laterally, with the exception of expatriates and rarespecies, with close ties to today’s North Atlantic biogeographic region. Climate related change in theAO is thus manifest at significantly differing time scales. Throughout �90% of the Pleistocene the AOhas existed in glacial mode, with narrow continental shelves, greatly restricted river inflow, thicker andperhaps immobile sea ice, and total blockage of exchange with the Pacific Ocean. During the Holocene,on shorter time scales of 1000–100 years, significant changes in high latitude climate are tied to changesin temperature and perhaps moisture delivery patterns. The Arctic also experiences significantmulti-decadal variability; however, the pace of change over the past three decades has been withoutprecedent. Within the basin interior the ice is now thinner and less compact, and thus more responsiveto wind stress (forcing and mixing). Concurrent with sea ice melt and increased river flow, the accumu-lation of fresh water and the stratification have increased, thus constraining vertical nutrient flux affect-ing phytoplankton size distributions, limiting primary production in parts of the basins now and likely inthe future, and increasing vulnerability to acidification. In addition, sea ice is now retreating on an annualbasis past the shelf break, exposing basin waters directly to sunlight and wind forcing. Thus, upwellingfavorable winds (generally from east to west) can now directly and efficiently drive shelf-break upwel-ling, and draw nutrients from subsurface basin waters onto the shelf; at the same time upwelling favor-able winds will also create onshore pressure gradients over the slope and basin which will act to slow orblock the flow of waters in the ACBC, and thus alter advective pathways of both abiotic and biotic mate-rials. Given the opening of a new ocean to multiple user groups, we expect that the central AO will play anincreasing larger role both in the research and political arenas in the future, and we encourage pan-Arcticinternational collaboration over focus on territorial boundaries.

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1. Introduction

Following earlier descriptions we here refer to the panarcticdomain north of Bering Strait in the Pacific andGreenland-Scotland Ridge in the Atlantic as the Arctic Mediter-ranean; the Norwegian, Iceland and Greenland seas south of FramStrait and west of the Barents Sea comprise the Nordic Seas, andthe remainder, including both shelf and basin domains, is the Arc-tic Ocean (AO). We refer to the central basins at the AmerasianBasin (AB) and the Eurasian Basin (EB). Our view of this oceanhas changed over time. It has changed because we have increasedour understanding of the area through painstaking scientific mea-surements since the time of the early Arctic explorers, to today’sincreasingly diverse suite of measurements by icebreakers, icecamps, moorings, ice-tethered profilers and satellites (Fig. 1). Ithas also changed because the system itself is in a rapid state oftransition (ACIA, 2005; Wassmann and Lenton, 2012; Polyakovet al., 2013; Bhatt et al., 2014). And central to understanding theAO system now and into the future is to understand its least stud-ied component: the deep basins.

At the onset of his historic voyage of 1893–1896 Nansenbelieved the entire AO to be a shallow sea (Nansen, 1902). Andwhile Nansen discovered that a deep basin occupied the central

AO, its trans-arctic ridge systems were unsuspected until the mid-dle of the 20th century (cf. Fig. 2a) which initially were deducedfrom tide and temperature data (Fjeldstad, 1936; Worthington,1953). Modern charts (Fig. 2b) now reveal two main basins, theEurasian (EB) and Amerasian basins (AB), separated by the Lomo-nosov Ridge. The EB is further separated into the Nansen andAmundsen basins by the Gakkel Ridge, and the AB into theMakarov and Canada Basin by the Alpha–Mendeleyev Ridge.Though the job of accurately mapping the AO bathymetry is farfrom complete, we recognize the important fact that more thanhalf (53%) the area of the present day AO is continental shelf, andthe remainder is the slope/rise/basin/ridge complex; (Fig. 2c;Aagaard et al., 1985; Jakobsson et al., 2012). To simplify the com-plicated bathymetry of the Arctic Mediterranean, and to link phys-ical and morphometric properties with biological distributions andfunctions, we follow Carmack and Wassmann (2006) and dividethe Arctic Mediterranean into fundamental hydro-morphologicaldomains. These are the inflow shelves, the interior shelves, theoutflow shelves, the Riverine Coastal Domain (Carmack et al.,2015), the two main basins, and the ridge and borderland features(cf. Fig. 2d). This distinction is critical to the understanding of basindynamics and biogeography owing to the importance ofshelf-basin interactions and inter-domain advection.

Fig. 1. Bar graph illustrating the number of ice camp and ship-based expeditionstaking place in the basins of the Arctic Oceans over the past six decades.

Fig. 2. Evolution of understanding of the Arctic basin’s bathymetry and hydro-morphom1944); (b) most-recent IBCAO map of the Arctic Ocean (Jakobsson et al., 2012); (c) simtypology of the Arctic Mediterranean based on its hydro-morphological domains: AW iset al., 2015); AO is Arctic Outflows; and arrows denote component flow directions.

B.A. Bluhm et al. / Progress in Oceanography 139 (2015) 89–121 91

The large ratio of shelf to basin area is one of the truly uniquefeatures of the modern AO; it is, however, a signature that haschanged often and rapidly in geologic time. The Arctic Basins beganto develop with the AB first ca. 140–150 Myrs ago, and then the EBca. 50–70 Myrs ago (cf. Backman and Moran, 2009). The AO’s deepconnection to the Pacific closed ca. 80 Myrs ago (Marincovich et al.,1990), leaving Fram Strait (sill depth 2600 m) as its only deep con-nection to the global ocean. The AO became the ‘modern’ Arctic ca.4 Myrs ago with the closure of the Central American Seaway(which began to freshen the Pacific relative to the Atlantic)(Zauker et al., 1994; Haug et al., 2001; Hasumi, 2002) and theopening Beringia Seaway (which allows the fresh Pacific water toflow back into the Atlantic via the AO pathways) (Coachman andBarnes, 1961; Carmack and McLaughlin, 2011). Ice core recordsshow that during this time, the Earth’s climate system began toresonate with orbital cycles, leading perhaps to theglacial-interglacial periods extant today (cf. Fedorov et al., 2006).In glacial mode the AO is almost all deep basin (except for the Bar-ents Sea) and is fed only by inflow from the Atlantic. In contrast,

etry: (a) bathymetric map of the Arctic Ocean at the mid-20th Century (Shirshov,plified bathymetry and place names used in this paper; and (d) highly idealized

Atlantic Water; PW is Pacific Water; RCD is Riverine Coastal Domain (see Carmack


during interglacial mode, as exists today, the Arctic is about 50%shelf and 50% basin, and functions as an Atlantic-Pacific estuary(Stigebrandt, 1981; Carmack and Wassmann, 2006). The abrupttransition to the present day Holocene conditions � 10 K yearsago thus marks an almost complete shift in variables and driversgoverning the Arctic marine system. The point here is that themodern (Pleistocene) AO spends �90% of its ‘time’ in glacial mode,with narrow continental shelves, greatly restricted river inflow,thicker and perhaps immobile sea ice, total blockage of exchangewith the Pacific Ocean, and shelf-basin interactions relativelyunimportant. It is also likely that the ‘functions’ of its borderingmarginal seas vary in geologic time. For example, water mass mod-ifications that currently take place on shelves would be absent dur-ing glacial mode, and the deep convection that currently takesplace in the Greenland Sea may move north into the EB in warmtimes and south into the Norwegian Sea in colder times (cf.Aagaard and Carmack, 1994). In this paper we focus on the Holo-cene AO, but recognize this as a transient phase to which the AObasin ecosystems may still be adjusting.

A succinct and interdisciplinary overview of the status and cur-rent understanding of the Arctic Basin domain remains elusive, inpart because there had been insufficient time and effort to explore,model and comprehensively describe basin properties and corre-sponding biological populations before climate change began toblur and confuse our work-in-progress picture. This paper thusaddresses but a small part of this larger challenge by presentingan overview of physical and biological conditions in Arctic basinswith a persistent focus on the role of advection in shaping ecosys-tems. First, a brief summary of AO physiography affecting advec-tion and exchange is presented (Section 2). Next, a simpletypology of the large-scale circulation which defines the contigu-ous domains (c.f. Wassmann et al., 2015), the water mass assem-blies, the sea ice climatology and freshwater sources relevant tobiogeography is offered (Section 3). Then, patterns in fuel for thefood web, including nutrient supply and primary production aredescribed (Section 4), and connections to zooplankton and benthosdistributions are drawn (Section 5). After that, special topics affect-ing biodiversity and carbon flows are addressed (Section 6) fol-lowed by an overview of physical and biological changes nowunderway (Section 7). Finally, a perspective on ongoing and futurechange is given (Section 8).

2. Basins, ridges, sills and gateways – the ever-changinggeometry of the Arctic Basins

History: The bathymetry and physiography of the AO form thebackdrop for ocean circulation, the vertical partitioning of watermasses, and the life contained within them. The major milestonesin mapping the bathymetry of the basin were not reached until1948–1950 when Russian expeditions finally mapped the Lomono-sov, Gakkel and Mendeleev Ridges (cf. Frolov et al., 2005) and withthe discovery of the Alpha Ridge soon following in 1963. To thisday, gaps still exist in our knowledge of the basins’ bathymetryand details of sill depths. Activities under the UN Convention onLaw of the Seas (UNCLOS) have renewed the interest in precisionseafloor mapping to resolve questions on territorial claims, andmassive bathymetric mapping efforts with a substantial focus onthe basins and shelf breaks are underway (Mayer et al., 2010;Jakobsson et al., 2012) and in 2012 culminated in a new versionof the International Bathymetric Chart of the AO (Jakobssonet al., 2012). In the central Arctic, one of the major updates includesmuch higher resolution of canyons along the continental slopes(Jakobsson et al., 2012), areas highly relevant for basin-shelf inter-actions such as dense water drainage and upwelling events (e.g.Aagaard et al., 1981; Melling and Lewis, 1982; Williams et al.,

2006; Williams and Carmack, 2008; Tremblay et al., 2011;Pickart et al., 2013).

Gateways, Basins, Ridges and Sills: Four fundamental gatewaysare recognized: Fram Strait, the Barents Sea Opening, Bering Straitand the Davis Strait; through these passages flow most of theArctic/sub-Arctic exchanges. Critically important for the interac-tion with the world ocean, Fram Strait (sill depth 2600 m) betweenGreenland and Svalbard forms the only present deep-water con-nection between the high Arctic and the North Atlantic. Exchangeof deep water and organic matter between the Arctic Basins andGreenland and Norwegian Basins is in both directions (e.g.Fahrbach et al., 2001; Schauer et al., 2004). Generally, inflow ofwarm and saline Atlantic water into the Arctic is along the easternside as the West Spitsbergen Current and through the Barents Seaopening, and outflow southward is on the western side as the EastGreenland Current. Shallow and narrow Bering Strait on the Pacificside provides a needle eye for Pacific-origin water to enter the AOwith a generally northward flow direction (Woodgate and Aagaard,2005). After crossing the Chukchi Shelf and undergoing seasonallyvarying modifications, Pacific-origin water primarily exits theChukchi Sea shelf through Barrow Canyon, Central Channel andHerald Canyon (Pickart et al., 2005, 2010; Weingartner et al.,2005, 2013).

The discoveries of the under-sea ridge systems, current flowsand vertical structure of the Arctic Basin waters combined withthe history of primarily national expeditions has shaped our cur-rent perception of the central Arctic as a tale of two main basinsand four sub-basins. The massive ridges across the central basinnot only divide the sea floor, but provide boundaries for watermass exchange and dispersal of biota (but see Section 6.2). Occupy-ing roughly 50% of the Arctic area combined, the EB and AB areinfluenced by their intimate connection to the Atlantic and PacificOceans, respectively, and by the often broad shelves around them.The very different characteristics of the Atlantic and Pacific watermasses entering these basins as well as the different depths andwidths of the inflow areas, Fram Strait, the Barents Sea Openingand Bering Strait, set up the vertical and horizontal water massstructure in the basins.

The central basins are classified into continental slopes andrises leading to rather flat abyssal plains that include some high-lands, and into large ridges (Jakobsson et al., 2003). The massiveridges form physical barriers between the major Arcticsub-basins. The ultra-slow spreading Gakkel Ridge (Dick et al.,2003) extends the North Atlantic Mid-Ocean Ridge system intothe Eurasian Arctic where it separates the Nansen and Amundsenbasins. Within its graben system is the deepest known locationin the AO (5243 m; Jakobsson, 2002). The Nansen Basin reachesdepths of 4000 m while in the Amundsen Basin extensive depthsof over 4500 m form the deepest plain of the central Arctic, thePole Abyssal Plain. The Lomonosov Ridge (sill depth �1870 m near88.7�N, 156.0�E; see Björk et al., 2007), stretches continuously allthe way from the Greenland continental slope to the Russian shelfoff the New Siberian Islands separating the Eurasian from the AB(Jakobsson et al., 2004). Within the AB, the massive Alpha–Men-deleev Ridge system (sill depth �2600 m at Cooperation Gap),which occupies almost 8% of the entire AO, separates the CanadaBasin (maximum depth �3800 m) from the Makarov Basin (maxi-mum depth �4000 m). This ridge, discovered during the drift of Icestation Alpha, is generally less steep and much more fragmentedthat the other two ridges. Several other large bathymetric featuressuch as the Yermak Plateau north of Svalbard and the ChukchiBorderland in the Canada Basin add additional complexity to theNansen and Canada basins. The continental slopes and the ridgesplay a major role in steering ocean circulation (Aagaard, 1989)and the distribution of biological communities (Deubel, 2000;Kosobokova and Hirche, 2009) as we will explore below.

Fig. 3. Historical perspective on Arctic Ocean connectivity and modern ice cover changes: (a) map from the mid-19th century illustrating the then prevailing notion of acentral Arctic Ocean kept ice-free by inflowing ocean currents (from Overland et al., 2011); (b) highly idealized schematic of flows linking arctic and subtropical watersaccording to the requirement to materials from regions where there is excess to regions where there is deficit (from Sandström, 1919); (c) map showing present day estimatesof transports and average T and S properties of the inflows and outflows based on an inverse model applied to boundary fluxes in summer 2005 (Tsubouchi et al., 2012), andmass balances estimates at Fram Strait (�2.0 ± 2.7 Sv, 1997–2006, Schauer et al., 2008), at the Barents Sea Opening (2.0 Sv, 1997–2007, Smedsrud et al., 2010), at Bering Strait(0.8 ± 0.2 Sv, 1991–2004, Woodgate et al., 2005; Melling et al., 2008) and Davis Strait (�2.3 ± 0.7 Sv for 2004–2005; Curry et al., 2011); and (d) map comparing the ArcticOcean ice cover (September minimum) for the 1979–2000 mean versus 2012 and overlaid on ocean bathymetry.


3. Of roundabouts, intersections and elevators – disentanglingthe physical oceanography of the Arctic Basin

History: During the early debates of the 19th Century on thecharacter of the Arctic Basin, the curious notion of an ice free cen-tral Polar Sea maintained by the warm inflows from the Atlanticand Pacific oceans was proposed. For example, Captain Silas Bent(1872) hypothesized that the Gulf Stream and Kuroshio, ‘‘penetratethe AO and carry with them warmth enough not only to dissolve allthe ice and snow they encounter in their paths, but enough also tokeep an open sea about the Pole at all seasons of the year’’ (cf.Overland et al., 2011; Fig. 3a). Sandström (1919) later discussedthe climatological requirement for arctic/subtropical exchangeand wrote, ‘‘The effects of currents originating in physical changein the water is simply that of transporting water from the region

where such water abounds to where it is rare . . . the complimen-tary physical processes whereby water is added to or drawn froma certain layer adapt themselves so as to balance one another inrate’’ (Fig. 3b). While the expeditions of the late 1800s and early1900s, for example Nansen’s Fram expedition in 1893–1896(Nansen, 1902), provided proof against the open Polar Sea concept,they did result in early evidence of Atlantic inflows to the deepbasins and of the primary direction of ice drift across the ArcticBasin via Transpolar Drift (Rudels, 2012). These findings were laterto be confirmed by observations from Russian ice drift stations(called ‘‘North Pole’’) that began in 1937 (Shirshov and Fedorov,1938). The associated fluxes of mass, heat and salt by Atlanticand Pacific waters have recently been summarized under the ban-ner of the Arctic and Subarctic Ocean Fluxes program (Dicksonet al., 2007, 2008; Fig. 3c). While the notion of an ice free polar

Fig. 4. Horizontal maps of annual mean near-surface (20 m) (a) salinity and (b) density north of 20�N for the Northern Hemisphere illustrate the basic estuarine circulationforced by low salinity and low density waters entering from the Pacific and more saline and denser waters entering from the Atlantic. Data source: http://odv.awi.de/en/data/ocean/world_ocean_atlas_2009/.


sea has long been dismissed, strictly, it is interesting that it doesdraw attention to advective inputs, which, ironically, may beaffecting today’s pattern in sea ice retreat (Fig. 3d; Shimadaet al., 2005; Carmack and Melling, 2011; Polyakov et al., 2012).

Present data, discussed next, suggest a wind-driven ice coverand upper ocean, a strong halocline complex separating the upperocean from the underlying waters of Atlantic origin, a weak buteddy-rich interior circulation rimmed by a contiguous system oftopographically-guided boundary currents, called collectively theArctic Circumpolar Boundary Current (ACBC), flowing cyclonicallyalong the margins of the ocean basins, and a very slow exchangeof deep waters within the basins (e.g. Aagaard, 1989; Rudelset al., 1994; McLaughlin et al., 2002; Timmermans et al., 2003;Aksenov et al., 2011).

3.1. Large-scale circulation

The AO is made integral to the global ocean by exchange flowswith the subarctic Atlantic and Pacific oceans, with the Atlanticdominating in terms of mass and heat fluxes and the Pacific dom-inating in terms of impact of freshwater flux on the vertical strat-ification (Aagaard and Carmack, 1989; Dickson et al., 2008). Thisclimate-scale system, evident in maps of surface salinity, density(Fig. 4a and b) and dynamic topography (Fig. 5a), is linked to thenorthern hemisphere hydrological cycle (or freshwater loop), inwhich ‘‘zonal distillation’’ due to the transport of water vapor fromthe Atlantic to the Pacific by Trade Winds across the Isthmus ofPanama, and ‘‘meridional distillation’’ due to net evaporation inthe low latitudes and net precipitation and river run-off in the highlatitudes result in a subarctic Pacific that is fresher and less densethan its Atlantic counterpart (Fig. 4a and b) (Aagaard et al., 2006;Carmack, 2007; Carmack and McLaughlin, 2011). The resultingsteric height difference (Fig. 5a) drives Pacific waters through Ber-ing Strait (sill depth �50 m) and into the Canada Basin (Coachmanand Barnes, 1961) where it encounters and overflows AtlanticWater that has entered the AO via both Fram Strait (sill depth�2600 m) and the Barents Sea (sill depth �400 m). Both inflowsencounter lighter surface waters freshened by river inflows andsea ice melt. This establishes the ‘‘estuarine circulation’’ of theAO with counter flows of warm, saline Atlantic waters belowcooler and fresher Pacific waters. Low salinity Arctic waters andsea ice exiting the AO via Fram Strait and the Canadian Archipelago

(mainly Nares Strait) are exported into the convective gyres of thesubarctic North Atlantic and then join the global thermohaline cir-culation which carries low-salinity waters back to the low lati-tudes, and therefore contributes to closing the global freshwaterloop (cf. Carmack and McLaughlin, 2011; for summary).

For the purpose of understanding biogeographical distributionswithin the AO, and to serve as a backdrop for an ecology of advec-tion (c.f. Wassmann et al., 2015), it is useful to recognize four basiclarge-scale (with L > 1000 km) circulation systems; these are: (1)the large scale wind-driven circulation of ice and the upper ocean(0 m to 50 or 150 m, depending on location) which forces thecyclonic Trans-Polar Drift from Siberia to the Fram Strait and theanticyclonic Beaufort Gyre in the southern Canada Basin(Fig. 5a); (2) the circulation of waters that comprise the haloclinecomplex, composed largely of waters of Pacific and Atlantic originthat are modified by freeze/thaw processes during passage overthe Bering/Chukchi and Barents/Siberian shelves, respectively(Jones and Anderson, 1986; Aksenov et al., 2011); (3) thetopographically-trapped Arctic Circumpolar Boundary Current(ACBC) which carries AW cyclonically around the boundaries ofthe entire suite of basins (Aagaard, 1989; Aksenov et al., 2011;Rudels et al., 2013; Fig. 5c); and (4) the very slow exchange ofAO deep waters which form initially in the Greenland Sea, enterthrough Fram Strait, and spread within the basin interior sequen-tially to the Nansen, Amundsen, Makarov and Canada basins(Macdonald et al., 1993; Schlosser et al., 1997; Rudels et al.,2012; Fig. 5d). Exchange of the latter from basin to basin is laterallyconstrained by the sequence of sills, but is aided in vertical motionby geothermal heating (Timmermans et al., 2003; Carmack et al.,2012). A summary of mid-water sources, pathways and associatedfronts are shown in Fig. 5d. Details of this circulation typology arediscussed next.

Surface Circulation: Much of what we know of surface circula-tion patterns stems from early observation collected from Russianice stations (cf. Coachman and Aagard, 1974; Frolov et al., 2005)and was deduced from observations of sea ice drift, starting withNansen (1902) and culminating in the International Arctic BuoyProgram (IABP) (Rigor et al., 2002; Pfirman et al., 2004). Sea iceand water masses lying above the cyclonically spreading Atlanticwaters (including the ACBC) are primarily wind-driven. Two basicsystems are manifest: (1) the cyclonic (Atlantic) Arctic bounded tothe east by the Siberian shelves and to the west by the Trans-Polar

http://odv.awi.de/en/data/ocean/world_ocean_atlas_2009/

http://odv.awi.de/en/data/ocean/world_ocean_atlas_2009/

Fig. 5. Schematic representations of Arctic Ocean circulation: (a) Surface circulation of the Arctic Ocean as shown by dynamic topography (20/400 dbar) (World OceanDatabase 2013, (b) summary of mid-water halocline sources, flows and associated fronts (blue shows Pacific-origin waters, maroon shows Atlantic-origin waters, thickmaroon line depicts the front between them) (after McLaughlin et al., 1996); (c) schematic representation of the Arctic Circumpolar Boundary Current system (ACBC) derivedfrom Atlantic water inflows (after Aksenov et al., 2011; Rudels et al., 2013); and (d) schematic representation of deep water exchange (Aagaard et al., 1985). BG is the BeaufortGyre, BSB is the Barents Sea Branch, FSB is the Fram Strait Branch, GG is the Greenland Gyre, NAC is the Norwegian-Atlantic Current, NCC is the Norwegian Coastal Current,TPD is the Transpolar Drift.


Drift and (2) the anticyclonic Beaufort Gyre over the Canada Basin(Fig. 5a). The cyclonic flow over the EB and some of the MakarovBasin defines a fundamentally divergent gyre, while the anticy-clonic flow of the Beaufort Gyre defines a convergent system(Fig. 5a); an important distinction in terms of regional stratificationand nutrient availability to the euphotic zone.

In recent years our understanding has been advanced by the useof satellite observations (sea surface height and bottom pressure)and much expanded in situ observations by ships, submarines, air-craft, ice camps, moorings and ice-tethered profilers (Morisonet al., 2012). From this has come a much deeper appreciation ofthe degree of variance from long-term mean in both velocity andin the spreading pathways of water masses. Proshutinsky andJohnson (1997) used a modeling approach to demonstrate twomodes of AO circulation according to patterns in atmospheric forc-ings. It was later suggested that such variability could influence thestorage and release of freshwater within the Beaufort Gyre into theconvective gyres of the Nordic Seas downstream (Proshutinskyet al., 2002). Subsequently, a regional-scale time series carried

out in the Canada Basin since 2003 has revealed a trend and largeyear-to-year variability in freshwater storage and strength of sur-face circulation within the Beaufort Gyre (McLaughlin et al.,2009; Proshutinsky et al., 2009; Krishfield et al., 2014).

Kwok et al. (2013) applied satellite data collected over a 28-yearperiod (1982–2009) to summarize the variance and shifts in deca-dal drift patterns in ice drift velocity and compared these to geos-trophic flow fields using the polarity of the Arctic Oscillation as abackdrop for atmospheric forcing. Mean circulation speeds overthe basin are of order 2–4 cm/s, with stronger flows along thesouthern edge of the Beaufort Gyre, in the Transpolar Drift, andespecially in the Fram Strait Outflow. These authors further reporta net strengthening of the Beaufort Gyre and the Transpolar Drift,especially during the last decade, with over 90% of the AO display-ing a positive trend in drift speed, and a decline in multiyear sea icecoverage. Importantly, they note the spatially averaged trendsfrom 2001–2009 in drift speeds (winter �+24%/decade, summer�+18%/decade) are not explained by the much smaller trends inwind speeds (winter �+1.5%/decade, summer: ��3.4%/decade),


with positive trends in drift speed in regions with reduced multi-year sea ice coverage. The increased responsiveness of ice drift togeostrophic wind is consistent with a thinner and weaker seasonalice cover and suggests large-scale changes in the air-ice-oceanmomentum balance related to reductions in ice strength andconcentration.

Halocline waters: Formation and subsequent circulation of halo-cline waters within the AO basins remains poorly known, and islikely due to a combination of mechanisms. Aagaard et al. (1981)and Melling and Lewis (1982) discussed the importance of seaice formation and brine drainage in modifying shelf-derivedwaters, and Killworth and Smith (1984) demonstrated the impor-tance of Pacific inflows to halocline formation. Rudels et al.(1996) proposed that wintertime convection in the Nansen Basinfollowed by summertime capping by ice melt and outflow of freshshelf water contributed to halocline formation. Kikuchi et al.(2004) further discussed the formation of convectively-formedwaters in the EB and its spatial and temporal variability, notingthat large-scale advance and retreat into the Amundsen Basinappeared to coincide with ocean circulation and frontal shifts asso-ciated with increased cyclonic circulation in the atmosphere. Jonesand Anderson (1986) used chemical distributions to distinguishbetween what they termed the upper (Pacific-derived) and lower(Atlantic-derived) halocline components. Bauch et al. (2014) usehydrochemical and oxygen isotope data do document analong-slope front between shelf, slope and central EB waters inthe Laptev Sea, and conclude that halocline waters above the slopeare derived from upstream advection. McLaughlin et al. (1996)argued that a front, the Atlantic-Pacific halocline front, spans theArctic basins where the Atlantic-derived waters subduct belowthe Pacific-derived waters (Fig. 5b, also Karcher et al., 2012). Theyfurther suggest that this front roughly aligns with and shiftsbetween the Lomonosov and Alpha–Mendeleev Ridge on decadaltime scales.

The inflow of PW is likewise complicated, with branches devel-oping south of Bering Strait from Anadyr Coastal Current waters,central Bering Shelf waters, and the Alaska Coastal Current waters,and continuing north across the Chukchi Sea via Herald, Hanna andBarrow canyons (Weingartner et al., 1998, 2013; Pickart et al.,2005, 2010). Due to the considerable widths of the Bering andChukchi shelves, and thus the long crossing time, these inflowsare modified en route to become the denser, winter and lighter,summer varieties of PW, with each branch undergoing differinggeochemical transformations (Nishino et al., 2013). Upon entryinto the AB, the tendency for PW to circulate in the cyclonic senseof the ACBC is opposed by the anticyclonic, wind-driven BeaufortGyre (cf. Shimada et al., 2006). Using modeling results, Aksenovet al. (2011) argued that cyclonic flow along the Alaska shelf andupper slope in the AB is forced by steric sea level differencesbetween the Pacific and Arctic. The confluence of Pacific waterswith halocline source waters draining from the Siberian shelvesresults in the Atlantic/Pacific Halocline front, separating the twodomains (Fig. 5c).

The Arctic Circumpolar Boundary Current: The input of AW to theAO has been studied for over a century but, because of its complex-ity and variability, pathways and fluxes are still subject to hugeuncertainties; a recent census of fluxes through the main gatewaysis given by Beszczynska-Möller et al. (2011). Between 8 and 9 Sv(1 Sv = 106 m3 s�1) enter the Nordic Seas over the Greenland Scot-land Ridge (sill depth �800 m) and of this roughly half continueson to the AO. Of the AW continuing north, about half enters theAO via Fram Strait as the Fram Strait Branch (FSB) and subductsbelow Arctic Surface waters (ASW) north of Svalbard. The otherhalf first crosses the Barents and westernmost Kara seas, subductsalong the Atlantic Polar Front, continues across the eastern BarentsSea and then enters the St. Anna Trough (Rudels et al., 2013). Here,

it is strongly modified by mixing with local Barents Sea waters andthe continuing eastward flow of the FSB to become the Barents SeaBranch (BSB) (Rudels et al., 2012, 2013; Dmitrenko et al., 2014,Fig. 5b). BSB water, having a broader density range than FSBwaters, both interleaves laterally and subducts below the continu-ing FSB. A third water mass, formed locally on the eastern Barentsand western Kara seas, drains into the basin through St. AnnaTrough (Aksenov et al., 2011). Subsequently, the three branchesbecome the ACBC and continue cyclonically around the basinperimeter, with bifurcations occurring where ridge and slopetopographies intersect. There is still debate as to the transportsof AW into and out of the AO, with a total amount or 4–8 Sv beinggenerally accepted (cf. Dickson et al., 2008).

Current meter measurements obtained during the Nansen andAmundsen Basins Observational System (NABOS) show the ACBCto flow along this pathway: from >20 cm s�1 in Fram Strait to15–20 cm s�1 northeast of Svalbard to 6–9 cm s�1 north of FranzJosef Land to 3–5 cm s�1 at the junctures of the Gakkel and Lomo-nosov ridges (I. Polyakov, pers. comm.). Note that slope currentstend to be strongest where the slope is steep (Isachsen et al.,2003). Aksenov et al. (2011) modeled the ACBC and demonstratedthat transports along the AO margins were forced by the jointeffects of buoyancy loss and non-local winds which created highpressure upstream in the Barents Sea. Spall (2007) applied aneddy-resolving numerical model and noted that lateral eddy fluxesfrom the boundary and vertical diffusion on the interior wereimportant drivers.

Deep Water: The pathways and rates of spreading of AO deepwaters are poorly known. In general, there is direct deep-waterexchange between the EB and the Norwegian and Greenland seasvia Fram Strait (sill depth �2600 m). From there the flow isthought to proceed from the Nansen Basin to the Amundsen Basinto the Makarov Basin and finally to the Canada Basin (MacDonaldet al., 1993; Schlosser et al., 1997) (Fig. 5c). From the AB there mustbe return flow back to the EB, Nordic Sea and North Atlantic(Aagaard et al., 1985). Processes of recirculation add complexityto the above circulation scheme, see Rudels et al. (2013) for furtherdiscussion. In terms of lateral advection of biotic material, it islikely that flows are strongest through gaps in ridges and withslope currents, especially where the slope is steepest (Isachsenet al., 2003).

The overall motion of deep water within the basins below silldepth is sluggish. The residence times of Arctic basin was deter-mined using 14C by Schlosser et al. (1997) who demonstrated alarge 14C gradient between the EB and AB, showing that the Lomo-nosov Ridge is a barrier to direct deep-water exchange between thetwo basins. They calculated the mean isolation age – the time thathas elapsed after a water parcel leaves the surface having acquiredits initial concentration via exchange with the atmosphere - of theEB bottom water (below 2500 m) to be about 250 years. (This iso-lation age is distinct from the average time a water parcel spendsin a particular deep basin.) The deep waters of the AB (below2500 m) are older than those of the EB, with an isolation age of450 years. Schlosser et al. (1997) found no significant horizontalor vertical gradients in 14C with in the AB (Makarov and Canadabasins) below 2250 m. Thus the deep waters on the North Ameri-can side of the Lomonosov Ridge are either presently not beingventilated (Macdonald and Carmack, 1991; Macdonald et al.,1993; Aagaard and Carmack, 1994), or are being ventilated slowlywith continuous renewal by shelf water (by freezing and brinerejection on the shelves) or influxes from the adjacent EB(Aagaard et al., 1985; Östlund et al., 1987; Jones et al., 1995;Rudels et al., 2000). Sinking brine plumes, if they do exist (see dis-cussions in Aagaard et al., 1985 and Rudels et al., 2013), may pro-vide a transport mechanism to carry organic material and biota todepth.


More rapid flows are expected along basin and ridge slopes, andthrough narrow gaps in the ridges. For example, the LomonosovRidge separates the EB and AB with average depths below sea levelbetween 1000 and 1400 m, but with a narrow gap of sill depth1870 m near 88�200N, 148�E. (Björk et al., 2007). This gap leadsfrom the Makarov Basin side of Intra Basin, a sub-basin atop theLomonosov Ridge. Timmermans et al. (2005) apply hydraulic the-ory to estimate a flow over the Lomonosov Ridge of0.25 � 106 m3 s�1, but Timmermans and Garrett (2006) speculatethat flows of this magnitude are not reaching fully to the bottom.Björk et al. (2007), however, present hydrographic data giving evi-dence that flow is actually from the Makarov to the AmundsenBasin, and subsequently interflowing at intermediate depths(1700–2000 m) as a return flow from the Canadian Basin to theEB and the Nordic Seas. Note that this direction of the water over-flow is opposite to that previously proposed by Jones et al. (1995)and Timmermans et al. (2005) and discussed as intermittent flowby de Middag et al. (2009). Geologic observations (Björk et al.,2007) and modeling results (Isachsen et al., 2003) are consistentwith strong current activity at the ridge crest, thus aiding plank-tonic dispersion.

Vertical motions may also play an important role in dispersionof biota. For example, an outstanding feature of the AO deep wateris its near homogeneity over thicknesses exceeding 1200 m withinthe Canada Basin and �800 in the Amundsen Basin, and capped bya temperature minimum (Timmermans et al., 2003; Björk andWinsor, 2006; Carmack et al., 2012). In the Canada Basin, for exam-ple, the deep waters vary in potential temperature (h) by less than0.001 �C between �2700 m and the bottom, a feature thatTimmermans et al. (2003) and Carmack et al. (2012) ascribe togeothermal heating and vertical convection. Salinity in the CanadaBasin increases with depth to �2700 m, but, like h, is nearly con-stant below this depth, at S = 34.957 psu. Björk and Winsor(2006) reported similar, near-homogenous deep waters in the EB,which they also attributed to geothermal heating. Using atime-series of deep water properties, Carmack et al. (2012) foundthat Canada Basin deep waters below �2700 m warmed at a rateof 0.0004 �C yr�1 between 1993 and 2010. This rate is slightly lessthan that to be expected from the reported geothermal heat flux(�50 mW m�2; cf. Langseth et al., 1990). Using this heat flux theyestimated a vertical velocity scale of � of 0.8 mm s�1, and thus forthe 1000 m-thick Canada Basin Deep Water the time-scale for con-vection is calculated to be �15 days. This value suggests rapid ver-tical mixing of the deep waters, particularly in relation to its longisolation age. The deep waters found above the lower continentalmargin of the deep basin maintain higher temperatures than thosein the basin interior, consistent with geothermal heat distributedthrough a shallower water column.

Lateral, inter-basin exchange of water along isopycnal surfacesat or near sill depths may effectively disperse biota. Carmack

Table 1Comparison between water mass properties of and river inflow into the Eurasian and Am

Eurasian basin Am

T/S of Polar mixed layer �1.7 to 4 �C/33.8–34.8 �1.Halocline properties One halocline/dS = 2 StepT/S of Atlantic layer 1–3 �C/34.8–34.9 0.6–T/S of deep water �0.9 �C/34.92 �0.Isolation age of bottom water 250 450Primary production (g C m�2 yr�1 �10–15 �1–Largest rivers (Mean discharge >100 km3 yr�1) Ob, Lena, Yenisey, Pechora,

Kolyma, Severnaya DvinaMa

River discharge (8 largest rivers) 65% 21%Sediment load (8 largest rivers) 17% 73%Atlantic/Pacific inflow (Sv) 5–8 �1

et al. (2012), for example, call attention to a deep temperatureminimum (hmin) layer that overlies the deep water, and that it isalso warming at approximately the same rate, suggesting thatsome geothermal heat escapes vertically through amulti-stepped, �300-m-thick deep transitional layer; double diffu-sive convection and thermobaric instabilities were suggested aspossible mechanisms governing this vertical heat transfer. Theseauthors conclude that the hmin layer is maintained by exchangewith the Makarov Basin, likely over the Alpha–Mendeleyev sillnear Cooperation Gap.

3.2. Water mass structure and distribution

Within the basin domain four basic water masses (according toorigin) exist; these are in order of increasing density: (1) The Arcticsurface waters ensemble (ASW), (2) Pacific-origin waters (PW), (3)Atlantic-origin waters (AW) and (4) Arctic deep waters. Increasedvertical gradients in salinity, or haloclines, separate these basicwater types (McLaughlin et al., 1996; Table 1). Representative ver-tical profiles of salinity (S) and potential temperature (h) for each ofthe four sub-basins illustrate the important point thatsalt-stratification increases several fold above the path of AtlanticWater flow from Fram Strait to the Canada Basin, owing to the pro-gressive accumulation of freshwater from various sources en route(Fig. 6). Of importance to biogeography and a consequence ofadvection, two basic water mass assemblies are observed (the termwater mass assembly refers to the vertical stacking of watermasses within the basins). The difference between them is theabsence or presence of PW sandwiched between ASW above andAW below. We refer to these domains as the Pacific Arctic andAtlantic Arctic, and the boundary between them as theAtlantic/Pacific halocline front (McLaughlin et al., 1996). This front(cf. Fig. 5d) is characterized by strongly sloping isohalines separat-ing the strongly and weakly stratified Pacific and Atlantic domains,and by the disappearance of the PW temperature maximum(Fig. 7). The key biophysical difference is that the Pacific Arctic ismuch more strongly stratified and is richer in nutrients, particu-larly silicate (Codispoti et al., 2005; Fig. 8). Both domains havesalt-stratification that constrains the vertical transfer of nutrientsto the surface layer (euphotic zone), thus leading to their charac-teristic oligotrophic state, particularly in the Pacific Arctic wherethe winter reset of surface layer nutrients by haline convection isvirtually absent (Codispoti et al., 2013). This difference is furtherdemonstrated by Laney et al. (2014) who used ice-tethered profil-ers equipped with bio-optical sensors to compare seasonal pat-terns in tow regions of the basins. They show that within theBeaufort Gyre region (Pacific Arctic) a deep chlorophyll maximumdevelops in summer, while in the Transpolar Drift System (AtlanticArctic) chlorophyll values are higher in the near-surface layer.

erasian Basins. Halocline values are typical of summer conditions.

erasian basin Reference

7 to 4 �C/28.0–34.4 JOISwise/dS up to 10 JOIS0.8 �C/34.80–34.85 McLaughlin et al. (2009)

5 �C/34.95 Björk and Winsor (2006) and Carmack et al. (2012)Macdonald et al. (1993) and Schlosser et al. (1997)

5 Codispoti et al. (2013)ckenzie, Yukon (indirect) Holmes et al. (2002)

Holmes et al. (2002)Holmes et al. (2002)Fahrbach et al. (2001), Schauer et al. (2008)and Woodgate et al. (2005, 2012)

Fig. 6. Representative profiles of (a) potential temperature (h), (b) salinity (S) and (c) h/S correlations for the four sub-basins of the Arctic Ocean: green is Nansen Basin, red isAmundsen Basin; yellow is Makarov Basin and blue is Canada Basin. Water mass assemblies are distinguished by the absence of presence of Pacific Waters sandwiched withinthe halocline. Note that the depth axis is plotted as square root of depth to better define the upper ocean where detail is greatest. Data from the JOIS program 2005–04 andfrom Polarstern ARKXII-1.


Surface waters of the AO are comprised of a base of either Atlan-tic or Pacific origin waters (depending on basin location) anddiluted by river inputs, ice melt and net precipitation(Yamamoto-Kawai et al., 2009). Far more river water is suppliedto the EB than to the AB (Holmes et al., 2002); however, the resi-dence time of these waters is relatively short, of order two years,as it is carried quickly from the basin into Fram Strait by the cyclo-nic flow of the EB and the Trans-Polar Drift (Anderson et al., 1989).In contrast, the accumulation of river water is especially pro-nounced in the Canada Basin, where the convergent winds of theatmospheric Beaufort High accumulate low salinity waters of bothNorth American and Siberia within the anticyclonic Beaufort Gyre,making this gyre the most strongly stratified component of the AO(Aagaard and Carmack, 1989; Proshutinsky et al., 2009). The result-ing heterogeneity in the distribution of freshwater components isreflected in maps of surface salinity (Fig. 4a).

Pacific Waters enter from the Bering Sea through Bering Strait,transit the Chukchi Sea and enter the deep basins along three mainbranches: one associated with Barrow Canyon; one east of HannahShoals; and one following Herald Canyon (Weingartner et al., 1998,2005; Aagaard et al., 2006). Inflowing waters include nutrient-richwater from the Gulf of Anadyr and fresher, lower nutrient AlaskanCoastal Cater (Walsh et al., 1989). While crossing the wide ChukchiShelf, is modified seasonally by biological production, heatexchange, ice formation and melting, and interaction with

sediments. Geochemical changes that occur here, associated withde-nitrification, impact the global nutrient cycle (Yamamoto-Kawaiet al., 2006).

Seasonal modification on the Chukchi Shelf produces two basicforms of Pacific-origin water that comprise the upper and middlehalocline of upper layer waters within the AB (Coachman andBarnes, 1961; Fig. 6c and 7). The upper halocline, due to summermodification and referred to as Pacific Summer Water, is character-ized by a local temperature maximum, between 31 < S < 32, inter-mediate nutrient and high oxygen concentrations. The middlehalocline, due to winter modification and referred to as PacificWinter Water, is characterized by a temperature minimum nearS = 33.1, high nutrient and lower oxygen concentrations.Coachman et al. (1975) observed the presence of two temperaturemaxima and associated the fresher one with Alaska Coastal Waterand the more saline one with Pacific Summer Water. Likewise,Shimada et al. (2001) observed two shallow temperature maxima,one at 31 < S < 32 and located east of the Chukchi Plateau and oneat 32 < S < 33 found west of the Chukchi Plateau. Steele et al.(2004) proposed that the pathways of Alaska Coastal Water andBering Sea Shelf Water are tied to variations in the Arctic Oscilla-tion Index (AOI), such that when the AOI is positive the PacificSummer Water crosses the Chukchi Plateau and follows the Trans-polar Drift and when the AOI is negative it also recirculates in theCanada Basin. An important consequence of this variability is its

Fig. 7. Section of salinity (S) colored by potential temperature (h) for an XCTD section crossing the Amerasian Basin; inset shows station locations. Note that the Atlantic/Pacific Halocline Front is evident in the steeply-sloping isohalines in the Canada Basin leading up to the Alpha-Mendeleev Ridge, and by the warmer Pacific varieties ofhalocline waters extant within the Canada Basin (from Kikuchi, Itoh, Eert and Williams, pers. comm.). Black vertical lines in bottom figure indicate locations of XCTDdeployments.


impact on the storage of freshwater in the Canada Basin, and, uponrelease back into the North Atlantic, affect convection in the sub-arctic gyres.

Atlantic Waters, as noted above, enter the Arctic basins as eitherthe warmer and more saline FSB or the cooler, fresher BSB, and iscarried basin-wide by the ACBC (Anderson et al., 1989, 1994;Rudels et al., 1994, 1999; Swift et al., 1997; Schauer et al., 1997,2002). Lateral spreading along near-isopycnal surfaces by thermo-haline intrusions also advect properties into the interior of both theEB (Perkin and Lewis, 1984; Rudels et al., 1994, 1999) and AB(Carmack et al., 1995; McLaughlin et al., 2004, 2009; Woodgateet al., 2007). Carmack et al. (1997) showed that such intrusionsare double diffusive structures in h and S, aligned or nested inh–S space across the Makarov, Amundsen and Nansen basins, andalso aligned in h–S space with those observed by Perkin andLewis (1984). They found intrusive layers to be 40–50 m thickand noted that their density increased as they spread laterally intothe basin interior. Because of their apparent robustness over suchdistance and time, Carmack et al. (1998) suggested that the layersmay be ‘self-organized’ and ‘self-propelled’ by the conversion ofpotential to kinetic energy by salt flux convergence, contractionon mixing or both. The presence of such features is allowed bythe weak ambient turbulence characteristics of the ice-coveredAO (c.f. Padman, 1995).

3.3. Sea ice

Unprecedented changes in sea ice have taken place in the pastdecades including a reduction in sea-ice extent (Cavalieri andParkinson, 2012) and thickness (Kwok et al., 2009), and a declinein the amount of multi-year ice (Comiso, 2012). The satellite recordreveals that over the past three decades the average summer min-imum has decreased by 11% per decade. Since the loss of areal cov-erage of summer cover exceeds that of winter, the area of theseasonal sea ice zone has increased (Kinnard et al., 2008). Impor-

tantly, with regards to this paper, is the fact that ice retreat isnow exposing surface waters over the deep basins; in 2012approximately 40% of basin area (with depth >400 m) was ice free.While not yet quantified, the effects of increased wind and solarforcing on the upper layers of the deep basins are likely to besubstantial.

Wind and atmospheric thermodynamic forcing were the maincauses for the record Arctic sea ice retreat in summer2007 (e.g.,Perovich et al., 2008). However, AW heat was likely importantfor preconditioning Arctic sea ice by making it thinner over severalpreceding decades and thus contributing to the extreme retreat(Polyakov et al., 2010). A similar situation exists for incomingwarm Pacific Summer Water spreading off shelf and into the Beau-fort Gyre in the Canada Basin (Shimada et al., 2006). Another fea-ture of observed decline in sea ice extent is that the trend is notstrictly linear, but instead consists of a series of punctuatedchanges, e.g. in 1989, 1998 and 2007 (see Perovich et al., 2014for time series). One plausible explanation is that ice thicknessand strength is slowly and inexorably decreased, year by year, byinternal forcing related to increased heat advection by the atmo-sphere and ocean, and then ‘shocked’ into a new dynamical stateby external forcings in extreme years. From the perspective ofcomplex systems behavior, thermodynamic forcing by the atmo-sphere and ocean acts as the slow variable to reduce ice coverstrength and resilience to external forcing, so that when excep-tional wind patterns do occur, such as in 2007 (Stroeve et al.,2012; Arctic Council, 2013), ice is readily exported south to meltin the Nordic Seas.

The above noted changes in sea ice are biologically significantsince sea ice and snow play several key physical and biogeochem-ical roles in the Arctic marine system. From a heat budget perspec-tive, the presence of sea ice drastically alters albedo and insulatesthe underlying water column from extreme winter heat loss. Themechanical properties of sea ice (thickness, concentration, rough-ness, etc.) greatly affect the efficiency of momentum transfer from

Fig. 8. Horizontal maps showing distributions of nutrient concentrations: (a) surface nitrate; (b) nitrate at 200 m; (c) surface phosphate; (d) phosphate at 200 m; (e) surfacesilicate; and (f) silicate at 200 m. Data derived from the Hydrochemical Atlas, CARINA, Codispoti et al., 2013 and JOIS (http://catalog.data.gov/dataset/hydrochemical-atlas-of-the-arctic-ocean-nodc-accession-0044630, http://cdiac3.ornl.gov/waves/discrete/, http://www.nodc.noaa.gov/archive/arc0034/0072133/1.1/data/0-data/).


the wind to the underlying water, with a thinner and loosely con-solidated ice cover being more effective in driving ocean currents.Freezing of sea ice within the basins and subsequent local meltingis observed to increase stratification, while the export and meltingof ice in the adjacent North Atlantic is a significant component in

the AO’s freshwater budget (Aagaard and Carmack, 1989). Sea icemelt water also has differing geochemical properties from ASW,such as heavier d18O values which can be used as a tracer(Macdonald et al., 1999), and lower alkalinity, a property thatmay exacerbate ocean acidification (Yamamoto-Kawai et al.,

http://catalog.data.gov/dataset/hydrochemical-atlas-of-the-arctic-ocean-nodc-accession-0044630

http://catalog.data.gov/dataset/hydrochemical-atlas-of-the-arctic-ocean-nodc-accession-0044630

http://cdiac3.ornl.gov/waves/discrete/

http://www.nodc.noaa.gov/archive/arc0034/0072133/1.1/data/0-data/


2009). Biologically, sea ice is the site of ice algal growth and trans-port, a habitat for ice meiofauna, zooplankton and small fish, and aplatform on which marine mammals may travel, hunt and breed(Bluhm and Gradinger, 2008).

Two aspects of sea ice reduction are worth special attentionwith regard to the deep basins. First, prior to the 21st centurythe sea ice margin seldom retreated beyond the shelf-break, andthus basin waters were seldom exposed to sunlight and wind. In2012, in contrast, roughly 40% of the area of the deep basins wasexposed (Fig. 2d). This would increase solar radiation to basinwaters, enhance wind mixing and increase shelf-basin exchange(see Section 6.3). The full physical and ecological consequencesof this new normal state (Wood et al., 2013; Jeffries et al., 2013)are currently a matter of debate, and on a regional basis, primaryproduction may either increase (Arrigo et al., 2008, forlight-limited shelf areas) or decrease (McLaughlin and Carmack,2010, for nutrient impoverished basins; Section 7). A secondimpact is that the reduction of ice in summer now allows enhancedgas exchange with the atmosphere, notably CO2, and this has beenimplicated as a factor in increasing the acidity of AO surface watersover the basins (Yamamoto-Kawai et al., 2011).

3.4. River inflow and other freshwater in the basins

Of all global oceans the AO is the most riverine, covering only3% of the global ocean surface area but capturing 10% of globalriver runoff within its mediterranean boundaries (Carmack,2000). Further, river inputs are increasing with global warming(Peterson et al., 2002, 2006), albeit with disproportional increasein Eurasian river run-off (McClelland et al., 2006). River inputs tothe AO also play a role in biogeochemical processes. Rivers bringturbidity that can counteract enhancing effects of nutrients byblocking sunlight to primary producers and cloggingfilter-feeders (Syvitski et al., 1989). Direct nutrient inputs from riv-ers are relatively small compared to the advective inputs from theAtlantic and Pacific Oceans (Codispoti and Owens, 1975), and likelysupport only �10% of total NCP (Gordeev et al., 1996; Gordeev,2000; see also Section 4.1). Alkalinity values, however, are typicallylower in incoming rivers than in AO waters, and their mixing withand inclusion in surface waters acts in increase local vulnerabilityto ocean acidification (Yamamoto-Kawai et al., 2011).

Liquid fresh water is also supplied to the AO through inputsfrom sea ice melt and net precipitation and, relative to an appropri-ate reference salinity by the Norwegian Coastal Current andinflows from the Pacific Ocean through Bering Strait. Sinks of liquidfresh water include export through the Canadian Arctic Archipe-lago and the western Fram Strait, and export of sea ice. Each ofthe sources contributes uniquely to halocline formation and struc-ture within the system, and to a variety of circulation and mixingprocesses that affect biological distributions.

An Arctic freshwater budget for sources, sinks and storage wasfirst produced by Aagaard and Carmack (1989), and has been sub-sequently updated by Serreze et al. (2006), Dickson et al. (2007),Tsubouchi et al. (2012) and Haine et al. (2015). Most estimates offreshwater content are based on the choice of reference salinity(for discussion, see Carmack et al., 2008). Aagaard and Carmack(1989) and Serreze et al. (2006) used 34.8, the mean salinity ofthe AO, while Dickson et al. (2007) calculated content relative toa salinity of 35, the salinity of incoming AW. Using available histor-ical data, Serreze et al. (2006) calculated a net AO freshwaterexport of 9200 km3 yr�1 and an import of 8500 km3 yr�1, leavinga net imbalance of 700 km3 yr�1. Using a similar approach, butwith a different reference salinity and study area, Dickson et al.(2007) computed an export of about 9500 km3 yr�1. Tsubouchiet al. (2012) applied an inverse model to constrain flux estimatesfor volume, heat and freshwater around the AO boundary (waters

above 1000 m) for a 32-day period in summer 2005. They calcu-lated mean properties for water entering the Arctic to beh = 4.49 �C and S = 34.50, and for water leaving the Arctic, includingsea ice, h = 0.25 �C and S = 33.81 (Fig. 3c). They calculated a corre-sponding volume transport into the Arctic of 9.2 Sv, an export of9.3 Sv, and a net oceanic and sea ice freshwater flux of187 ± 48 mSv (Fig. 3c).

Regional circulation processes play a huge role in the distribu-tion and pathways of freshwater components within the basins.Initial river inflow at source creates a coastal-trapped RiverineCoastal Domain flowing from west to east and rimming much ofthe inner continental shelf (cf. Carmack and McLaughlin, 2000;Carmack et al., 2015, Fig. 2d). Upon exiting the shelves, for exampleduring upwelling favorable wind events, river water mixturesserve to freshen the surface waters within the basins. In general,wind-driven currents serve either to store fresh water, as is thecase for the convergent Beaufort Gyre, or to export fresh water,as is the case for the divergent Trans Polar Drift. Upon export intothe convective gyres of the subarctic North Atlantic the low salinityarctic outflow acts in poorly understood ways to influence convec-tion and the global overturning cell.

Because of its importance in determining stratification withinthe AO and its potential importance in governing deep convectionin the adjacent subarctic Atlantic, much effort has gone intoobserving and modeling variability in freshwater budget of theAO, both in terms of fluxes into and out of the basin and storage.The original estimate of Aagaard and Carmack (1989) was thatthe EB held 12.2 � 103 km3 of liquid fresh water while the AB heldthe larger volume of 45.8 � 103 km3, despite the fact that the sumtotal of river inflows to the EB is much larger than that to theAmerasian. The total freshwater content at this time, includingshelf domains and sea ice, was found to be near 100 � 103 km3,approximately equal to the freshwater stored globally in lakes.The larger volume in the AB was attributed to the storage of PWwithin the Beaufort Gyre. Subsequently, the retreat of the Paci-fic/Atlantic halocline front from the vicinity of the Lomonosov tothe Alpha–Mendeleyev Ridge in the late 1980s and early 1990s(McLaughlin et al., 1996; Morison et al., 1998) would have neces-sarily reduced this volume. Morison et al. (2006) argued for areturn of the AO to pre-1900s hydrography, but did not place thisstrictly in the context of freshwater storage. Proshutinsky et al.(2002) advanced a numerical and a conceptual model to purportthat storage and subsequent release of freshwater from the Beau-fort Gyre and export to the Nordic Seas could inhibit deep conven-tion. Since then, a number of studies have discussed interannualvariability in AO freshwater content as forced by winds, increasedriver inflow and sea ice melt. Proshutinsky et al. (2009) examinedyear to year variability in the southern Canada Basin and found anunprecedented increase in freshwater storage of �25% during thefirst decade of the 21st Century. McPhee et al. (2009) and Gileset al. (2012) also confirmed the increase in freshwater storagewithin the basin associated with wind-driven convergence in theBeaufort Gyre. Rabe et al. (2014) calculated a trend in freshwaterstorage between 1992 and 2012 for the entire AO (for bottomdepths >500) of 600 ± 300 km3 yr�1. A decrease in salinity madeup about 2/3 of the freshwater trend and a deepening of the upperlayer the remaining 1/3. Time and space variability in storage andpathways are discussed in papers by Newton et al. (2008), Morisonet al. (2012), and Korhonen et al. (2012).

In any discussions of the AO freshwater budget, its future stateand its role in biogeochemical processes, it is critically important toidentify the source of the fresh water, be it river, ice melt, directprecipitation or Pacific inflow, this is done through the judicioususe of geochemical tracers such as d18O, barium, N/P ratios andalkalinity (Carmack et al., 2008; Yamamoto-Kawai et al., 2008,2009). This is important biologically, because freshwater from


the Pacific comes with labile, marine-derived carbon while riversbring terrestrial hard-to-digest carbon into the basins (Brownet al., 2014); difference that have implications for the foodweb (Iken et al., 2010; Dunton et al., 2012). Also, each of thesesources will respond differently and at different rates to climatechange.

4. Fuel for the food web – Nutrients, carbon and production

4.1. Nutrients

The delivery of nutrients to the Arctic basins is complex, and thedynamics of source and sink poorly understood. The main sourceultimately derives from the multiple inflowing streams of Atlanticand Pacific waters, and to some extent, by inflowing rivers (cf.Holmes et al., 2012). Each source then follows a different pathwayand is modified en route by seasonal and regional biogeochemicalprocesses before subducting below the ASW and circulating withinthe basin. The basins themselves comprise an oligotrophic systemwith strong salt stratification, extensive ice cover, and low sunangle, all of which limit new production (Tremblay and Gagnon,2009; McLaughlin and Carmack, 2010; Codispoti et al., 2013;Varela et al., 2013). Strong salt stratification and damping of windmixing by ice cover (when present) constrain vertical mixing andthe attendant vertical supply of nutrients to the euphotic zone;hence the lateral advection of nutrients combined with episodicmixing events can be of disproportionate importance in maintain-ing production and in determining the size spectra of phytoplank-ton (Li et al., 2009). In the AB and parts of the EB a prominentsubsurface chlorophyll-a maximum layer lies immediately atopthe summer nutricline which, itself, is associated with the upperboundary of Pacific Summer Water (McLaughlin and Carmack,2010). In this sense, the supply of nutrients by inflowing subarcticwaters to the deep basins behaves in the manner of a ‘filling box’(cf. Killworth and Smith, 1984), with the subsurface interflow driv-ing vertical flow to supply surface outflow.

The distinction in nutrient concentrations between the Pacificand Atlantic water mass assemblies is evident in surface and200 m maps of NO3 (Fig. 8a and b), PO4 (Fig. 8c and d) and SiO4 con-centrations (Fig. 8e and f). We did not filter the data sources by sea-son, but most data were collected during the summer. Nitrate isnear-depleted in the surface waters of the anticyclonic AB whileresidual values (2–5 lmol l�1) persist in the cyclonic EB (Fig. 8a).This implies that primary production is nitrate limited in the AB,while in the EB it is constrained by either light, grazing or both(cf. Codispoti et al., 2013). Nitrate levels at 200 m are dominatedby Pacific water inputs (Fig. 8b). Surface values of phosphate inthe basins are generally above zero, ruling out phosphate as a lim-iting nutrient (Fig. 8c). Similarly, the 200 m levels of phosphate aremuch higher in the AB than in the EB, a signal of de-nitrification onthe Chukchi Shelf that is carried with the arctic outflow on into theNorth Atlantic where it plays a role in the global nitrogen cycle(Fig. 8d; Devol et al., 1997; Yamamoto-Kawai et al., 2006). Surfacevalues of silicate are relatively low in the incoming Atlantic watersand extending across the Barents Sea; they are also low in the cen-tral Beaufort Gyre, perhaps –we speculate – reflecting differentialsinking rates of diatoms in a domain of low f-ratio (Fig. 8e). Silicatelevels at 200 m are very high in the AB, reflecting Pacific waterinputs, and strong gradients are evident across the Atlantic/Pacifichalocline front (Fig. 8f). Silicate distributions also show that the ABsupplies outflow to both the Canadian Arctic Archipelago and Baf-fin Bay.

Differential nutrient concentrations in the two central basinsmay constrain their responses to seasonal processes. Some nitrate,albeit in low concentrations, remains present in surface waters of

the EB after the spring bloom in summer and may be availablefor ice marginal blooms into the autumn. In contrast, virtually nonutrients remain in the upper water column after the spring draw-down in the oligotrophic Beaufort Gyre centered over the AB,hence preventing marginal ice blooms from forming. In both basinsdelayed freeze-up may result in enhanced wind-mixing in fall thatcould deliver nutritients to the euphotic zone, stimulate a fallbloom, and thus increased production (Loeng et al., 2004; Ardynaet al., 2014).

Freshwater buoyancy fluxes are very large during summer overmost AO due to ice melt, river runoff, and the low salinity (partic-ularly during summer) of the inflowing Pacific Waters (Carmack,2000; Woodgate et al., 2006; Fig. 3c). These processes create rela-tively shallow mixed layers well into the central basins. Summermixed layers in the AO are often <10 m deep (e.g. Codispoti et al.,2005), and throughout most of this system mixed layer depths sel-dom exceed 50 m, even during winter (Steele and Boyd, 1998;Carmack, 2007). The strong stratification places an important con-straint on the nutrient supply to the photic zones within thebasins. The principal exception within the AO is the southern Nan-sen Basin adjacent to the Barents Sea (Codispoti et al., 2013).

Within the basin interior, and removed from shelf-break pro-cesses, the vertical supply of new nutrients to the surface (eupho-tic) zone is carried out by two main mechanisms: (1) halineconvection in winter during ice formation and growth, whichresets the summer (vegetative season) drawdown, and (2) by ver-tical mixing. Haline convection, the flux of which sets up and isthen depleted by the spring bloom, is severely constrained by saltstratification, especially in the AB (Codispoti et al., 2013; Varelaet al., 2013). The latter, which supports the deep chlorophyll max-imum (cf. Tremblay et al., 2009; McLaughlin and Carmack, 2010), isenhanced by current shear and internal wave dissipation, and ishighly variable in time and space. For example, higher vertical fluxvalues are to be expected above the shelf-break and submarineridges where currents are faster, and this increases nutrient supplyto the deep chlorophyll maximum. Tidal motions over rough andvariable topography may likewise generate vertically propagatinginternal waves that enhance vertical mixing within the halocline(Kulikov et al., 2010). Vertical mixing can also be enhanced in openwater by wind forcing in late summer and fall prior to freeze-up,potentially resulting in small fall blooms (cf. ACIA, 2005; Ardynaet al., 2014).

Inflowing rivers also supply nutrients and organic carbon, but inmuch smaller quantities than supplied by currents, and impactsare largely confined to their adjacent shelf regions (Holmes et al.,2012). In a recent biogeochemical budget study for the AO,Macdonald et al. (2010a,b) estimated the AO imports an equivalentof 42.5 kmol N s�1 from the Pacific and Atlantic oceans comparedto 2.5 kmol N s�1 from rivers (Macdonald et al., 2010a,b). Dissolvedorganic nitrogen loads from rivers to the AO can also potentiallycontribute to the dissolved inorganic nutrient pool throughphoto-oxidation or bacterial degradation (Holmes et al., 2012;Tank et al., 2012).

4.2. Primary production – pelagic and sea ice

As on the Arctic shelves, algal biomass in the basins results fromphytoplankton and ice algal production and there is much discus-sion about the magnitude and relative contributions of those twocomponents and their variability related to snow and ice cover(Gradinger, 2009; Leu et al., 2011, 2015; Matrai et al., 2013). Whilethe contribution of ice algal production to total primary productionhas been estimated at 4–25% on the Arctic shelves (Legendre et al.,1992), it can be >50% across the basins (Gosselin et al., 1997) andeven up to 90% in the Canada Basin (when including sub-ice upperwater column production; Matrai and Apollonio, 2013). In a


modeling study, the sea ice in the Arctic basins contributed asmuch as 18% to total ice algal biomass in the Arctic based on thesheer size of the area, while daily production rates and productionper unit area were among the lowest in the Arctic (Deal et al.,2011). The earlier ice algal bloom (mean production from Marchto May, cf. Jin et al., 2012) plays an important role as an early foodsource and fuel for reproduction of sympagic and some pelagictaxa (Runge and Ingram, 1988; Søreide et al., 2010; Leu et al.,2011), before the subsequent pelagic bloom occurs (typically fromMay to September; Jin et al., 2012). More recently, the formation,distribution and density of ice algal aggregates in and under seaice has been illuminated in the central Arctic, and these aggregateswere suggested to contribute significantly to vertical export ofice-derived carbon (Fernandez-Mendez et al., 2014; Katlein et al.,2014).

Generally, both sea ice algal and pelagic primary production inthe two basin domains are one or two orders of magnitude lessthat on adjacent continental shelves (Arrigo et al., 2008;Gradinger, 2009; Table 1). The underlying causes outlined aboveare nutrient limitation due to strong stratification and light limita-tion due to snow and ice cover and extreme sun angle (Sakshaug,2004). Severe nutrient limitation favors small phytoplankton (Liet al., 2009), e.g. flagellates over diatoms, a low f-ratio of new torecycled carbon fixation (Varela et al., 2013), and a low energy foodweb that favors jellyfish over fish and marine mammals (cf.Parsons and Lalli, 2002). In the AB the virtual absence of nutrientsin the surface waters combined with relatively high values in theunderlying PW leads to a ubiquitous deep chlorophyll maximumimmediately atop the interflowing PW (McLaughlin and Carmack,2010). For this reason, estimates of phytoplankton biomass basedon remote sensing data alone may underestimate productivity.

To explore net community production (NCP) patterns on apan-Arctic scale, Codispoti et al. (2013) estimated pelagic NCP fromseasonal changes in nutrient concentrations. They note that regio-nal heterogeneity in NCP is sufficiently large within the AO (�twoorders of magnitude) as to make meaningful comparisons. In theAB they found very low nitrate concentrations in the upper 50 m(�0 mmol m�3) even in winter and correspondingly low annualNCP (�1–5 g C m�2 yr�1), and concluded that nutrient limitationsuppresses NCP in this region. Low wintertime nitrate concentra-tions in the upper layers of the AB suggest that winter reset ofnutrients by convective mixing is vanishingly small, and thatNCP in these sub-regions may, therefore, be insensitive to changesin the ice and light regimes. In this area, new estimates suggestthat sympagic primary production may contribute an unusuallyhigh proportion, as much as 90%, to total basin net community pro-duction (Matrai and Apollonio, 2013). In the EB in contrast, pelagicNCP is >10 to �15 g C m�2 yr�1, and significant surface layer nutri-ent concentrations persist throughout summer (Fig. 8), suggestingthat light or grazing or both may limit NCP.

While productivity estimates are critically important for under-standing food web dynamics, knowledge of which species con-tribute to the production and their size, function, phenology, etc.is arguably as important. Indeed, on the pan-arctic scale, there isa great diversity of primary producers with 1874 known speciesof phytoplankton and 1027 species of ice algae documented sofar (Poulin et al., 2011). Most of these species are large cells(>20 lm), dominated by centric and pennate diatoms, anddinoflagellates. In the basins, taxonomic composition differsbetween surface and sub-surface chlorophyll maximum layers,although biomass was similar in that study (Coupel et al., 2012).The contribution of smaller phytoplankton and ice algae to totalmicroalgal diversity and their contributing to total phytoplanktonbiomass is only now beginning to be appreciated (Li et al., 2009;Collins and Deming, 2011; Lee et al., 2012).

5. Of pattern and process, scarcity and hotspots – distributionsin faunal biomass and food webs

Much of the carbon produced in the upper water column and inthe ice is consumed by sympagic and pelagic grazers in the upperlayers, and by omnivores and detritivores throughout the watercolumn and at the seafloor. As a result of low in situ productionin the basins and consumption of freshly produced carbon whileparticles are sinking, vertical carbon flux to the deep-sea floor iscomparatively low (Olli et al., 2006). Particles advected from pro-ductive shelves such as the Barents, Chukchi and Kara Sea shelvesand from turbidites add substantial amounts of carbon to the ver-tical supply (Grantz et al., 1996; Cooper et al., 1999; Soltwedel,2000). While the areas around the basins perimeter receive thisallochthonous input, biota in the central Arctic Basins away fromthe shelves do not. As a result, biomass tends to decrease withincreasing water depth and/or distance from the shelf both in thewater column and at the sea floor (Fig. 9; e.g. Clough et al., 1997;Bluhm et al., 2011a; Kosobokova and Hirche, 2000, 2009; Weiet al., 2010; Kosobokova, 2012).

5.1. Zooplankton

The few historical assessments of the zooplankton biomass inthe Arctic basins reported very low biomass with less than 0.2–3.0 g dry weight (DW) m�2 in the 0–1500 m layer, but were diffi-cult to compare due to methodological differences or incompletesampling of the water column (Minoda, 1967; Hopkins, 1969a,1969b; Pautzke, 1979; Kosobokova, 1982; Conover and Huntley,1991). Ashjian et al. (2003) suggested that the zooplankton stockmay have been significantly underestimated in many of these earlystudies, and subsequent studies undertaken in the 1990s and early2000s indeed reported greater zooplankton biomass and produc-tion (Wheeler et al., 1996; Mumm et al., 1998; Thibault et al.,1999; Kosobokova and Hirche, 2000; Ashjian et al., 2003;Hopcroft et al., 2005). Recent comprehensive assessments at >80locations scattered over the EB and AB and using consistent meth-ods estimated zooplankton biomass integrated over the entirewater column at ca. 2–24 g DW m�2 dry mass (Kosobokova andHirche, 2009; Kosobokova and Hopcroft, 2010; Kosobokova,2012). Regional variability is strongest within the EB, which isundoubtedly related to the circulation patterns. A prominent beltof elevated abundance and biomass is found along the Eurasiancontinental slope from north of Svalbard to where the LomonosovRidge meets the continental slope (Figs. 10 and 11a). The elevatedbiomass is clearly confined to the core of the Atlantic inflow (ACBC)(Kosobokova and Hirche, 2009) that suggests that the ACBCadvects plankton populations from the North Atlantic, and thatthey to a large extent remain within the zone affected by it(Wassmann et al., 2015). This also includes enhanced biomass inrecirculating branches of the Atlantic inflow along mid-oceanridges as the Nansen–Gakkel Ridge (Hirche and Mumm, 1992)and the Lomonosov Ridge (Kosobokova and Hirche, 2000). Biomassdrops sharply to a third or half of the slope values toward the deepbasins (Fig. 11a, Fig. 12; 2–3.8 g DW m�2). Elevated biomass at thebasin periphery roughly coincides with the marginal ice zone andseasonally ice-free waters (Kosobokova and Hirche, 2001, 2009)where primary and, subsequently, secondary production are highercompared to the until recently permanently ice covered and hencefood-limited basin centers (Sakshaug and Slagstad, 1991;Sakshaug, 2004; Ulsbo et al., 2014).

Almost everywhere in the Arctic basins vertical profiles of zoo-plankton abundance and biomass in summer are characterized bymaximum concentrations in the 0–50 m layer and a noticeabledecrease, by several orders of magnitude, with depths (Hopkins,

Fig. 9. Schematic sections illustrating the distributions of water masses and vertical biomass distributions along representative sections crossing the Arctic basins; insetshows the location of sections: (a) from the Bering Sea across the Lomonosov Ridge and into the Greenland Sea; (b) from the Bering Sea and through the Canada Basin andCanadian Arctic Archipelago into Baffin Bay, (c) from the Bering Sea across the Lomonosov Ridge and onto an interior shelf off Siberia. Colored circles represent averagebiomass distribution across all basins of mesozooplankton (red) and macrobenthos (blue) with water depth (modified from Kosobokova, 2012 and Bluhm et al., 2011b). AWAtlantic Water, DW Deep Water, HC Halocline, PW Pacific Water, SW Surface Water, SAT St. Anna Trough; white dotted line denotes the approximate upper boundary of thehomogeneous layer.

Fig. 10. Map showing overlay of expatriate zooplankton distributions with halocline circulation patterns and the Atlantic/Pacific halocline front (colors as in Fig. 5d). Onlyrecords deeper than 500 m are shown. Data from Kosobokova (2012).


1969a; Kosobokova, 1982; Ashjian et al., 2003; Kosobokova andHopcroft, 2010). The uppermost 0–50 m layer, where all freshlyproduced carbon is concentrated, contains 40–50% of overall bio-mass in summer (Kosobokova and Hopcroft, 2010; Kosobokova,2012; Fig. 9), as elsewhere in the World oceans. In principal, allvertical zooplankton distribution patterns are alike in both Arctic

basins (Fig. 12a and b), but the depth of the maximum dependson the seasonal state of the zooplankton community and the inter-play between populations of migrating species. Regional variabilityis related to the proximity of the Atlantic gateway area (Fig. 12a,profiles highlighted in blue), where advected populations of theAtlantic copepod Calanus finmarchicus cause a biomass maximum

Fig. 11. Maps showing (a) zooplankton biomass (g DW m�2) (modified from Kosobokova and Hirche, 2009; Kosobokova and Hopcroft, 2010) and (b) infauna (macrofauna)biomass (g C m�2) (modified from Bluhm et al., 2011b) for the two Arctic basins. Biomass is concentrated along the shelf breaks and in inflow areas. Only records deeper than500 m are shown.

Fig. 12. Vertical distribution of zooplankton biomass (mg m�3) in (a) the Eurasian and (b) Amerasian basins. Blue are stations close to pronounced AW or PW influence, redindicates stations north of 85�N, green shows the stations distributed elsewhere. Data from Kosobokova and Hirche (2009), Kosobokova and Hopcroft (2010), and Kosobokova(2012).


at mesopelagic depths confined to the core of Atlantic inflow.Similarly, locations in the AB close to the Pacific inflow (Fig. 12b,blue profiles) also show higher biomass at all depths.

Two major components contribute to the zooplankton biomassand production of the AO under present conditions, an autochtho-nous and an allochthonous community, and their comparison interms of species composition, community structure and life cycletraits helps explain regional biomass distribution patterns(Kosobokova and Hirche, 2009, see also Pomerleau et al., 2014for comparative community structure). The former consists oflocally reproducing species that maintain resident populations in

the AO including the basin-wide distributed key species, the largearctic calanoid copepods, Calanus hyperboreus and C. glacialis,Metridia longa and the chaetognath Eukrohnia hamata. Their localproduction forms the biomass base in all central basins and overtheir slopes. The allochthonous community, in contrast, consistsof plankton populations advected from the contiguous domains(cf. Wassmann et al., 2015), but are not able to reproduce in arcticwaters and becoming sterile expatriates there (Kosobokova andHirche, 2009; Kosobokova, 2012; Wassmann et al., 2015). The mostimportant contributors to allochthonous stock, the expatriatecopepods C. finmarchicus advected with Atlantic inflow, hardly


penetrate into the AB due to their short, one year long life spans.The AB receives hardly any allochthonous addition to locallyproduced biomass, because the majority of plankton advected fromthe North Pacific already dies off on the extended shelf of theChukchi Sea. Consequently, oceanic communities of the AB arecharacterized by significantly lower biomass compared to the EB(Fig. 10).

Although there are pronounced seasonal variations of zooplank-ton abundance and biomass in the surface water layer of the Arcticbasins (Kosobokova, 1982, 2012; Ashjian et al., 2003), the verticalzooplankton distribution pattern shown in Fig. 12 undergoes littlechanges during the annual cycle. Year-round observation in the ABrevealed that abundance and biomass per unit volume at depths donot increase substantially below 100 m even when key Arctic zoo-plankton species descend to overwintering depths, because theyare distributed over a wide depth range of several hundred meters(Geinrikh et al., 1983; Kosobokova, 1982, 1983, 2012; Ashjian et al.,2003; Darnis and Fortier, 2014). Density maxima remain withinthe upper 0–100 m layer throughout the year (Kosobokova, 1982,2012; Ashjian et al., 2003).

Unique to the Arctic basins, the lower margin of the epipelagiczone seems to be shallower than in the rest of World oceans whereit occupies the upper 200 m of the water column (Vinogradov,1970). An abrupt drop in zooplankton biomass below 100 m(Fig. 12) (Kosobokova, 2012), a noticeable upward shift of verticalranges of common mesopelagic species compared to elsewhere(‘polar emergence’, Kosobokova, 1989, 2012; Kosobokova andHopcroft, 2010), and a boundary between statistically distinct epi-pelagic and upper mesopelagic zooplankton communities at�100 m (Kosobokova et al., 2011) suggest that the epipelagic risesup to depths of about 100 m in the Arctic. This phenomenon isprobably related to a combination of limited light penetration intothe water column due to low sun angle, presence of sea ice withsnow cover, strong stratification of the upper water column, andoverall low primary production concentrating fresh food in a thinsurface water layer (Zenkevitch, 1963; Harding, 1966;Kosobokova, 1989, 2012).

The basin zooplankton fauna uses several mechanisms to copewith the food limitation. In the surface layers, ice-associated andupper water column pelagic crustaceans are closely tied to freshalgal production as reflected in their relatively low d15N ratios,indicators of low trophic level (Iken et al., 2005). The deeper plank-ters use different food sources, because a considerable part of thefresh organic matter is already converted into fecal material andmarine snow by ice-associated and surface-dwelling planktonicherbivores and recycled the microbial loop within the epipelagiczone before it sinks to depth (Olli et al., 2006). Consequently, car-nivory and omnivory/detritivory are prominent feeding modes inmeso- and bathypelagic Arctic zooplankton communities(Harding, 1974; Laakmann et al., 2009; Kosobokova et al., 2002,2011). A number of these deep planktonic carnivores and detriti-vores have modified gut passages as an adaptation for more effi-cient digestion and assimilation of nutritionally poor anddigestively resistant food. Sigma-shaped (e.g. in the copepodAetideopsis rostrata) or substantially widened guts (e.g. in the cope-pods Scaphocalanus acrocephalus and S. polaris) permit longerretention of food inside the gut (Kosobokova, unpublished). Oneof the most striking examples is the copepod Spinocalanus antarcti-cus which possesses a strongly elongated and looped midgut,enabling it to digest marine snow and organic coating oftiny-sized mineral particles melting out of the ice and sinkingthrough the water column (Kosobokova et al., 2002). Vertical nichepartitioning, whereby closely-related zooplankton species occupydifferent depth ranges to reduce resource competition, is anotheradaptation used by the AO deep-water plankters in their

resource-restricted environment (Kosobokova et al., 2007;Laakmann et al., 2009; Kosobokova and Hopcroft, 2010).

5.2. Benthos

As for zooplankton, abundance and biomass also decreases withwater depth at the Arctic (and global) deep-sea seafloor (Soltwedel,2000; MacDonald et al., 2010a,b). This decrease tends to be greaterfor benthic macrofauna (infauna) than for meiofauna (Wei et al.,2010). This trend has been interpreted as an average decrease inmetazoan size with increasing water depth (Klages et al., 2004;Rex and Etter, 2010) and as increased importance of smaller organ-isms with increasing water depth (Pfannkuche and Soltwedel,1998; Rex et al., 2006). Generally, the range of macrofaunal densi-ties and biomass (mostly below 4000 ind m�2 and <1 g C m�2,respectively, summarized by Klages et al., 2004 and Bluhm et al.,2011a; Fig. 11) fall within the lower end of values reported fromthe North Atlantic (Levin and Gooday, 2003). Meiofaunal densitiesin the central Arctic under ice-cover are either lower than in theglobal deep sea (Schewe, 2001) or on the same order of magnitudeas in other oligotrophic deep areas (<100 to >3000 ind 10 cm�2;Soltwedel, 2000; Vanreusel et al., 2000; Hoste et al., 2007;Górska et al., 2014). Despite the general trend of decreasing densityand biomass with depth, however, faunal densities and biomassmay vary substantially in areas of similar depths, depending onvertical in situ and advective carbon flux to the seafloor that isrelated to ice-edge effects and vicinity to inflow areas in the AO(Schewe and Soltwedel, 2003; Bluhm et al., 2005, see also Sec-tion 6.3). Klages et al. (2004) concluded cautiously, based on therelatively sparse biomass and oxygen consumption rates availablefor the Arctic deep-sea benthos, that consumption rates are gener-ally both low, though regionally variable with lowest rates in thebasins, and in agreement with those from other deep-sea areas,but that vertical organic matter supply may be higher thanexpected from calculations of vertical carbon flux from sedimenttraps.

The food web at the sea floor of both Arctic basins is still poorlydescribed. We do know, however, that it is characterized by a highdegree of reworking of organic material resulting in four to fivetrophic levels (excluding marine mammals; Iken et al., 2005;Bergmann et al., 2009; van Oevelen et al., 2011). Benthic procary-otes and, in the larger size classes, deposit feeders play a major rolein carbon recycling as reflected in the dominance of procaryotes ina carbon flow model in the EB (van Oevelen et al., 2011) and inenriched d15N signatures of macrofaunal biomass-dominantdeposit feeders and those of their predators and scavengers inthe AB (Iken et al., 2005). Typical for the deep sea, suspension feed-ers are less common in the abyss of the AB (Bluhm et al., 2005;MacDonald et al., 2010a,b) because of extremely small currents(Timmermans et al., 2010), although this feeding type is moreprevalent in the more rapidly moving deep waters in Fram Strait(Bergmann et al., 2009, 2011). Interestingly (and somewhat con-trary to the above concept that smaller organisms become moreimportant with depth), some larger members of the benthic com-munity in both Arctic Basins are apparently capable of quickly andefficiently ingesting and utilizing fresh material when it doesbecome available, which has been long been documented fromother deep-sea areas (e.g. Billett et al., 2001). The rapid respondto food pulses kilometers under the Arctic sea surface was recentlyconfirmed from observations of holothurian (sea cucumber) con-sumers of fresh ice algal matter at the seafloor (Boetius et al.,2013) and low d15N ratios (indicating low trophic level) in someadditional benthic taxa (Iken et al., 2005, Iken and Bluhm unpub-lished). In distant locations in both Arctic basins, the samesurface-feeding mobile holothurians, Kolga hyalina, congregated


in comparatively high densities (maximum >30 ind m�2) to consumefresh organic matter from the sediment surface (MacDonald et al.,2010a,b; Boetius et al., 2013) as known for their relatives in thedeep North Atlantic (Billett and Hansen, 1982). In contrast, thenutrition of nematodes, the density-dominant metazoan meio-fauna in Arctic deep-sea sediments, is still unclear, although micro-bial production are a suspected food source based on correlation ofmicrobial activity with meiofauna densities (Hoste et al., 2007), buta lack of uptake of microbial isotopic markers by nematodes couldnot confirm the trophic relationship (Guilini et al., 2010).

Fresh algae, Melosira arctica, in K. hyalina guts were apparentlyabundant and nutritious enough to spur reproduction shortly afterthe food pulse in the EB (Boetius et al., 2013). These recent obser-vations suggest that the paradigm change from the 1980s–1990sthat many deep-sea taxa in fact reproduce seasonally (triggeredby food pulses) rather than continuously (due to a perceived lackof a zeitgeber at the time; e.g. Witte, 1996) also applies to the Arc-tic basins. This also holds true for the omnivorous and carnivorousdeep-sea zooplankters in the AB and EB. Although able to repro-duce at low rates all year round, they show clear peaks ofegg-laying activity confined to the Arctic summer followed bymaximum abundance of offspring (Kosobokova, 2012).

6. Who is where and why? Faunal distribution in light of historyand physiography

6.1. Pacific, Atlantic or Arctic? – Origin, biodiversity and biogeographyof the Arctic Basins’ biota

The Arctic basin’s patterns in diversity and biogeography areclosely tied to the geological history and semi-isolation of the areaas well as to the connections with the Pacific and Atlantic Oceans(Golikov and Scarlato, 1989). Originally an embayment of theNorth Pacific, the Arctic deep sea was influenced by Pacific faunauntil �80 million years ago when the deep-water connectionclosed (Marincovich et al., 1990). Exchange with the deep Atlanticbegan �40 million years ago, coinciding with a strong cooling per-iod (Savin et al., 1975).

Three different mechanisms are conceivable for the coloniza-tion of the Arctic basins by benthic taxa (Mironov et al., 2013).These authors summarized that some eurybathic benthic shelf taxain the Atlantic Arctic supposedly expanded their range into theArctic deep sea during Pleiostocene glaciations, while other shelffaunas were eradicated. Those taxa that migrated to depths areby some authors considered the ancestral fauna of some of today’sArctic deep-sea fauna (Gurjanova, 1985; Nesis, 1984). Speciesinventories of Arctic shelves (Piepenburg et al., 2011) and basins(Bluhm et al., 2011a) showed that indeed �60% of the macro-and megabenthic deep-sea species (>500 m) are shared with theArctic shelves. Mironov et al. (2013) argue that the second andthird lines of argument are that the Arctic deep-sea fauna is dom-inated by and originates from fauna common in other areas of theworld’s deep sea, either the North Atlantic or the North Pacific.Multivariate analysis of Eurasian, Amerasian and Greenland-Iceland-Norwegian Seas polychaete species distributions supportedthis second line of argument in that it revealed a strong (recent)Atlantic influence and the absence of modern Pacific fauna(Bluhm et al., 2011a). Also, Mironov et al.’s (2013) synthesis docu-mented that of over 100 species from 92 genera in 8 classes otherthan polychaetes, over half of all genera were deep-sea specialists,with closer ties to the Norwegian and Greenland Seas (albeit inpart as a transition area) than the North Pacific on the species level.Those authors caution, however, that all possible source regions ofthe Arctic Basins fauna are of multiple origins themselves, compli-cating the disentanglement of biogeographic origins.

Similarly, the deep-sea zooplankton fauna is also believed tooriginate primarily from other areas of the deep World’s oceans.Almost half of all resident zooplankton species of the AO have widedistribution ranges with the majority of those being cosmopolitan(25%), bipolar (9%), or species in common with the North Atlanticand the North Pacific (10%) (Kosobokova et al., 2011). The modernNorth Atlantic fauna contributes 25%, although it was long believedthat at least half of the zooplankton species in the Arctic Oceanwere of North Atlantic origin due to free deep water exchangethrough Fram Strait (Brodsky, 1956; Brodsky and Pavshtiks,1977; Grainger, 1989). Taxa in common with the North Pacificare very few (2%) and their influence on the modern metazoanplanktonic fauna of the AO can be largely neglected (Kosobokova,2012). As with the benthic fauna, faunal exchange between theNorth Atlantic and the AO through Fram Strait has taken placefor the last 18 million years, whereas the shallow Bering Straitand Chukchi Sea prevented penetration of mid- and deep-waternon-migrating plankton species from the North Pacific into theArctic for the past 80 million years. In contrast to the benthos,however, hardly any shelf planktonic taxa expand their range intothe Arctic deep sea, or show submergence into the abyssal. Theunderlying reasons are the large differences in ecological require-ments of epipelagic neritic (shelf) inhabitants relative to the condi-tions available at great depths. Therefore, only two of the threehypotheses of colonization of the Arctic basins by deep-waterfauna suggested by Mironov et al. (2013) could explain the originof the Arctic deep-sea zooplankton.

A recent faunal species inventory of Arctic Basins’ benthosyielded more than 1100 taxa recorded deeper than 500 m andnorth of 80�N in Fram Strait (Bluhm et al., 2011a). Relative to anearlier inventory (Sirenko, 2001), over 400 new taxa were added,primarily in the speciose groups of round worms (Nematoda), bris-tle worms (Annelida) and crustaceans. Taxa with very few specieswere also species poor globally. As elsewhere in the global deepsea, the Arctic Basins have a high fraction of rare and endemic spe-cies, although the estimates (50–80% endemism rate perVinogradova, 1997 and Mironov et al., 2013) may be artificiallyhigh because of the relatively low sampling effort. Several dozennew species have been discovered in the Arctic Basins in recentyears (e.g. Rogacheva, 2007; Gagaev, 2008) although species dis-covery rates are by far lower than those in the Antarctic deep ocean(Brandt et al., 2007), because of lower sampling effort in the Arctic,and high faunistic connectivity to the North Atlantic. Dozens ofnew distribution locations, however, are recorded by anydeep-sea expedition, and are thought to reflect the poor knowledgeof the fauna rather than actual recent species range extensions (e.g.MacDonald et al., 2010a; Sharma and Bluhm, 2011).

The largest recent metazoan plankton fauna inventory is basedon collections from 134 locations where sampling was conductedfrom the surface to the bottom or 3000 m depth during recent ice-breaker expeditions and older Russian drifting ice stations(Kosobokova et al., 2011). The inventory of 174 species from eighttaxa (Cnidaria, Ctenophora, Mollusca, Annelida, Nemertea, Crus-tacea, Chaetognatha, and Larvacea) is now assumed to be nearlycomplete with the exception of the deepest water layers, whereboth unrecorded and new species continue to be found(Kosobokova et al., 2011). The number of mesozooplankton speciesin the basins is about half of that in the entire AO including itsshelves (the latter ca. 370 species). The epipelagic basin fauna islow in diversity (Kosobokova et al., 2011) and consists of a fewubiquitous forms and seasonally migrating species endemic to arc-tic waters, such as the copepods Calanus hyperboreus, C. glacialis,and Metridia longa. Of the 174 species recorded in the Basins, 38species are recent additions to earlier lists (Sirenko et al., 1996;Kosobokova et al., 1998; Sirenko, 2001), and 18 of those wererecently described (Markhaseva, 1998, 2002; Markhaseva and

Fig. 13. Map showing the distribution of some (a) pelagic and (b) benthic taxa occurring in the Eurasian and Amerasian basins. Taxa are identified in the legend. Thedistributions suggest that the Lomonosov Ridge does not serve as a distribution barrier. Note that for many other species, even fewer distribution records exist for the Arcticbasins preventing a general conclusion on distribution patterns. Data from Bluhm et al. (2011b) and Kosobokova et al. (2011).


Kosobokova, 1998, 2001; Raskoff et al., 2010; Andronov andKosobokova, 2011) or are currently under description. Crustaceansstrongly dominate zooplankton species number (70%), and cope-pods are the most diverse group among them. Only 14% of all zoo-plankton species in the basins are endemic to the Arctic(Kosobokova, 2012) with one-third of them ice-associated andthe remaining two-thirds bathypelagic with a clear preferencefor depths below 1000 m (Kosobokova et al., 2011). The evolutionof these deep-water endemics is without doubt related to thegeological history of the AO (Kosobokova et al., 2011).

6.2. The role of the ridges – steering currents, but no barrier forpan-Arctic Basin species dispersal?

The prominent ridge systems (Section 2) do not act as strictzoogeographical barriers for zooplankton or benthic invertebratefauna between the EB and AB (Koltun, 1964; Deubel, 2000;Bluhm et al., 2011a; Kosobokova et al., 2011). The samedeep-water and endemic zooplankton species are found on bothsides of the Lomonosov Ridge (Fig. 13a) and the vertical structureof epipelagic, mesopelagic and bathypelagic communities are con-sistent between basins (Kosobokova et al., 2011). Some deep waterbenthic species and similar communities are also found on bothsides of the Lomonosov Ridge (Bluhm et al., 2011a,b; Mironovet al., 2013; Fig. 13b), although the high rate of endemic and rarespecies at the Arctic Basins’ seafloor limits generalization of thisfinding (MacDonald et al., 2010a; Mironov et al., 2013). The currentpatterns suggest regular and recent exchange of at least some ofthe deep fauna across the ridge, and find some support in recentinsights of the circulation patterns in the basins. For example,deep-water flow was recently documented between the Makarovand Amundsen basins through the deepest sill (�1870 m) in theLomonosov Ridge allowing water and biotic exchange betweenAB and EB deep waters (Björk et al., 2007, 2010). The distributionof the bathypelagic zooplankton assemblage below roughly1000–1500 m also suggests physical and biological exchangesacross this ridge sill are possible (Kosobokova et al., 2011),although exact boundaries in zooplankton community structurecannot be drawn because of limited vertical resolution in the sam-pling technique. If arctic deep-water zooplankton taxa undergo

large vertical migrations, this could also serve as a mechanismfor exchange between basins if horizontal exchange of fauna acrossthe ridge is or recently was occuring. As noted above, free exchangefrom the EB to the AB is likely down to the sill depth of 2400 m(Timmermans et al., 2005) and from the Makarov to the CanadaBasin through Cooperation Gap in the Alpaha/Mendeleyev Ridgewith sill depth of 2400 m (Carmack et al., 2012).

It is unclear, however, exactly how benthic fauna disperseacross the Arctic deep sea and over the ridges. Dispersal distanceestimates are sparse for any region, but are thought to be a func-tion of larval period, current velocities, habitat, larval biology andadult genetic makeup (Grantham et al., 2003; Palumbi, 2003 andreferences therein). On the one hand, conditions in the deep Arcticbasins facilitate long-distance dispersal, because larval life spanstend to be longer at low versus high temperatures (Fetzer andArndt, 2008) and in soft bottom versus other habitat types(>30 d; Grantham et al., 2003). In addition, predation pressure inthe deep water column is likely low given the low zooplanktonand nekton densities in deep layers (Raskoff et al., 2010;Kosobokova et al., 2011). On the other hand, dispersal is probablyconstrained by low horizontal flow rates in near-bottom Arcticdeep water (<1 cm s�1; see Timmermans et al., 2010, but also obvi-ous from high densities of animal tracks despite low faunal densi-ties, Macdonald et al. (2010a)), and pelagic life stages are generallythought to be less common in high latitudes than elsewhere(Thorson, 1950; although challenged by Fetzer and Arndt, 2008and others). We speculate that larval dispersal in the Arctic basinsmight be aided by upward convective transport by geothermalheating that has been estimated at 1 mm/sec (Carmack et al.,2012), followed by lateral transport over ridge sills within theoverlying water (at �1 cm s�1; Timmermans et al., 2010) for hori-zontal dispersal in flow rates higher than in waters below silldepth, especially near steep topography. Based on these assump-tions, pelagic larvae could rise to the overlying layer in a few days,disperse laterally across the Lomonosov Ridge if they started in theAmundsen Basin, for example, and settle in the Makarov Basinwithin a month. Taxa without pelagic larvae such as amphipodsand isopods would be expected to disperse much more slowly,and perhaps have restricted distributions as found for, e.g. Antarc-tic isopods and ostractods (Brandt et al., 2007), although we found


no clear evidence for the latter in the literature. A comprehensivecompilation of dispersal stages and distances of Arctic fauna hasyet to be done.

In the upper water column, halocline, and Atlantic layers, Atlan-tic and Pacific expatriate zooplankton taxa are advected with theirrespective waters either into the EB or AB (Kosobokova and Hirche,2009; Kosobokova and Hopcroft, 2010; Kosobokova, 2012; Nelsonet al., 2009, 2014; Wassmann et al., 2015). In addition, inhabitantsof the shelves may be advected offshore as neritic expatriates(Kosobokova, 2012; Wassmann et al., 2015). Although abundancesof expatriates in the AB are much lower than in the EB, they doresult in differences in the species composition between basinsin surface waters and midwater layers. Recent Pacific benthicspecies, in contrast, are virtually limited to the shelves, in particu-lar the Chukchi and Beaufort Sea shelves (Bilyard and Carey, 1980).

6.3. Two-way traffic: shelf to basin and basin to shelf interactions

The Arctic Basins are intricately influenced by their surroundingshelves, and the shelves are influenced by the deep central Arctic,especially along the margins. Above, we have already discussed therole of the shallow PW inflow for the water mass structure andstratification and the role of AW inflow for biogeography, biodiver-sity and faunal biomass patterns. In this section, we focus on phys-ically driven shelf-basin interactions and their biologicalconsequences.

6.3.1. Shelf to basinHere, we summarize key shelf-basin interactions, recognizing

that other mechanisms are concurrently operating. Particularlyimportant is the role of ice formation and brine drainage overshelves in the maintenance of the halocline complex in the basins(Aagaard et al., 1981; Melling and Lewis, 1982). While the subarc-tic North Pacific and North Atlantic are sources of halocline waters,the physical and biogeochemical composition of the water massesfrom these areas undergo substantial changes during transit of theinflow shelves before entering the basins (Jones and Anderson,1986). These authors proposed that brine production during seaice formation creates saline bottom waters which inhibit verticalmixing with overlying fresh riverine water entrainment ofre-mineralized nutrients into the photic layer. The constituentnutrients are then advected horizontally into the halocline layerof the deep basins (c.f. Popova et al., 2013). Similarly, Nitishinskyet al. (2007) and Dmitrenko et al. (2011) have documentedenhanced near-bottom nutrient concentrations in the Laptev Sea.The influence of this ‘continental shelf pump’ on halocline watershas been further investigated by Anderson et al. (2010) whoshowed that the Chukchi and East Siberian seas supply an excessof carbon and nutrients to the Canada Basin; and by Alkire et al.(2010) who used the tracer NO (NO = 9 � [NO�3 ] + [O2]) to demon-strate a Siberian shelf influence in the Makarov Basin.

Dense water drainage from shelves occurs preferentiallythrough cross-shelf troughs such as Barrow Canyon and the St.Anna Trough (Garrison and Becker, 1976; Mountain et al., 1976;Dmitrenko et al., 2011). The general trend of decreasing faunal bio-mass with increasing water depth from the shelf toward the basinis interrupted in such areas, because carbon sources are concen-trated and channeled from the productive shelves into the less pro-ductive basins. The western Beaufort Sea slope downstream of PWdrainage through the Barrow Canyon is one such location. Biomassof Arctic cod, snow crab and certain epifaunal invertebrates arehigher along the shelf break of the Western Beaufort Sea than onthe adjacent shelf (Crawford et al., 2011; Logerwell et al., 2011;Parker-Stetter et al., 2011; Ravelo et al., 2015). Belugas , whoseprey include Arctic cod (Frost and Lowry, 1984; Loseto et al.,2009; Quakenbush et al., 2015), tend to use this area on the upper

continental slope of the Beaufort Sea that coincides with the com-paratively warm Atlantic layer (Suydam et al., 2005; Citta et al.,2013). The occurrence of belugas along the western Beaufort slopeis also correlated with a well-developed Alaska Coastal Currentproducing strong frontal features (Stafford et al., 2013) and theseauthors suggest this mechanism would provide enhanced foragingopportunities.

Mesoscale eddy formation provides another mechanism for offshelf transport of material properties and biota. Baroclinic eddieshave long been recognized to populate the halocline of basin inte-riors (Hunkins, 1974; Newton et al., 1974); an early census sug-gests that eddies comprise up to one quarter of the surface of thesouthern Canada Basin (Manley and Hunkins, 1985). Morerecently, Zhao et al. (2014) examined ice-tethered profiler datadeployed between 2004 and 2013 to carry out a census of mesos-cale eddies within the AO halocline in the basins. They docu-mented 127 eddies, 95% of which were anticyclonic, the majorityof which had anomalously cold cores and were observed in theBeaufort Gyre (AB eddies) and the Transpolar Drift (EB eddies).Such eddies typically have horizontal length scales on the orderof 10–20 km and horizontal (azimuthal) velocities of order 10–25 cm s�1 (Timmermans et al., 2010) and their formation has beenattributed to instabilities related to current/topography interaction(D’Asaro, 1988), frontal zone processes (Timmermans et al., 2008)and shelf break jets (Pickart et al., 2005, 2013; Spall et al., 2008).Offshore transport of shelf-origin waters by eddies of Pacific originin the upper halocline are associated with elevated concentrationsof nutrients, organic carbon, and suspended particles (Mathis et al.,2007; Watanabe et al., 2012; Pickart et al., 2013). These authorsattributed these features to derive from the boundary currentalong the edge of the Chukchi Shelf. O’Brien et al. (2011) examinedsediment trap data from the slope off the Canadian Beaufort Shelfand argued that eddy phenomena also played a major role in trans-porting sediments from shelf to basin. For the Canada Basin, eddiesin three different depth domains were recognized: (1) shallowupper halocline eddies centered at �80 m), (2) lower haloclineeddies (�200 m) and (3) deep eddies (�1200 m) (Carpenter andTimmermans, 2012). Eddies observed in the deep waters at depthsbetween 200 and 2000 m of both the AB (Swift et al., 1997;Carpenter and Timmermans, 2012) and EB (Schauer et al., 2002;Aagaard et al., 2008) had water mass properties derived from adja-cent shelves. Velocities within deep eddy cores were found torange from 2–25 cm s�1, and thus appear to have the potential totransport material properties within and below the Atlantic layerand re-suspend particles where Atlantic water abuts the continen-tal slope (Carpenter and Timmermans, 2012).

Sea ice, through its drift patterns, also transports large amountsof particles from rivers and the shelves into the central AO (Eickenet al., 2005). This transport mechanism is in fact thought to havecritically shaped the sedimentation regime in the AO in the geolog-ical record (e.g. Nørgaard-Pedersen et al., 1998). Particle types arecomprised not only of sediments and terrigenous carbon that are inpart region-specific in terms of their composition and can, there-fore, be used as tracers (Dethleff et al., 2000), but also of fresh algalmaterial that provides an allochthonous food source for basinfauna (Boetius et al., 2013).

6.3.2. Basin to shelfWind forcing is the principal mechanism driving basin waters

onto the shelf, either via upwelling favorable winds (generallyeasterlies) bringing water from depth onto the shelf, or via down-welling favorable winds (generally westerlies) driving surfacewaters onto the shelf. In the arctic wind forcing may take placein regions that are ice covered, partially ice covered or ice free(see Wadams, 2000, for discussion). Clearly, however, as sea icecontinues to retreat, more and more of the shelf break domain is


exposed annually to wind-forced, shelf-break upwelling (Carmackand Chapman, 2003). In the Canada Basin this trend has the poten-tial to increase the volume of nutrient-rich Pacific Water onto theshelf, and potentially increase primary productivity (Tremblayet al., 2011; Spall et al., 2014). Williams et al. (2006, 2008) exam-ined the joint roles of wind-forcing and bathymetry (e.g. canyons,isobaths divergence) on upwelling and cross shelf transport.

The efficiency of upwelling in driving water exchange across theshelf-slope boundary is strongly affected by topography and thepresence of canyons, headlands and convergent isobaths. Forexample, Williams et al. (2006) examine shelf-break exchange inMackenzie Canyon, a cross shelf canyon in the Beaufort Sea shelfforced by wind- and ice-driven ocean surface-stresses. They notethat while the canyon is approximately 400 m deep and 60 kmwide, it is only 2–3 times the baroclinic Rossby radius at its mouth,and thus behaves as a dynamically ‘narrow’ canyon. Upwellingevents are associated with wind in the ice-free summer seasonand with ice motion in winter (see also Pickart et al., 2013), butice motion does not necessarily correspond to wind-stress becauseinternal ice stresses differentially block downwelling-causing icemotion (e.g. it is easier to move ice offshore than onshore). Thisasymmetry between upwelling and downwelling flow, combinedwith the regional predominance of upwelling-causing ice motion,show that Mackenzie Trough is an effective conduit for transport-ing deeper, nutrient-rich water onto the shelf (Williams et al.,2006). In the Mackenzie Trough Carmack and Kulikov (1998)observed vertical isopleth displacements associated with AtlanticWater exceeding 400 m and, upon cessation of wind forcing, col-lapsing and then relaxing back into the basin to generate a Kelvinwave and its associated velocity field propagating eastward alongthe slope.

In addition to the large canyons such as Mackenzie, St. Anna andBarrow canyons, there are hundreds of shallow submarine valleysthat are relict channels of lower Pleistocene sea levels during theice ages. Many such channels are shallow and do not extend asfar as the shelf edge, so that their role in cross-shelf exchange inresponse to upwelling favorable winds is unclear and likely differ-ent from case to case. For example, Williams et al. (2008) examinedflow in Kugmallit Valley on the Canadian Beaufort Shelf and indeedfound cross shelf flows associated with upwelling and down-welling wind forcing. However, because the valley ends out beforereaching the shelf break, upwelling found within the valley musttransport water derived from across the shelf rather than acrossthe shelf break.

Upwelling and cross-shelf exchange can also be greatlyenhanced by isobath convergence/divergence. For example, hydro-graphic and remote sensing data from shelf waters north of CapeBathurst show consistent evidence of upwelling events (Williamsand Carmack, 2008). These authors propose that this enhancedupwelling is forced by the vorticity adjustment of the along-shelfflow to the isobath divergence at the cape. Benthic samples nearthe cape show enhanced numbers and diversity of organismswhich suggest that nutrients brought to the surface by upwellingallow enhanced primary production in the region that ultimatelyfeeds the benthos (Conlan et al., 2013).

The biological consequences of accelerated seasonal ice retreatand enhanced upwelling of basin halocline waters onto adjacentshelves are dramatic because increases in new production requirean upward shift in the delivery of nitrate to the euphotic zone(Tremblay and Gagnon, 2009). To demonstrate this, Tremblayet al. (2011) quantified the effects of shelf-break upwelling in theCanadian Beaufort Sea where repeated episodes of ice retreat andupwelling favourable winds resulted in 2 to 6-fold increases in pri-mary production. They also explain that interpretations of changein arctic productivity must distinguish between new production,which yields a net increase in plant biomass due to increases in

nitrogen delivery, and regenerated production, which is supportedby the recycling of biological products (see also Codispoti et al.,2013).

Upwelling and downwelling favorable winds will also affect thealong-slope and off-slope transport of the ACBC. Lien et al. (2013)show events in which the relative strengths of the FSB and BSB areaffected by wind-forced Ekman transport in the northern BarentsSea, such that the decrease in sea surface height resulting fromupwelling winds induces a cyclonic circulation anomaly alongthe slope which, in turn, weakens the FSB flowing along the basinslope. If these events got more common we speculate that we mayobserve a general, pan-arctic wide pattern involving weakening ofproperty transport by the ACBC, with concurrent decreased deliv-ery of allochthonous material to far-field portions of the basins,counter-posed by increased production of autochthonous materialby shelf-break upwelling, a hypothesis which would need rigoroustesting.

7. Of halocline structure, primary production and vertical flux –change underway?!

No matter what is used as a clock, rapid change has become thehallmark of modern AO research, and this, in turn, is manifest atsignificant and differing time scales. As noted before, throughout�90% of the Pleistocene the AO exists in glacial mode, with narrowcontinental shelves, greatly restricted river inflow, thicker and per-haps immobile sea ice, and total blockage of exchange with thePacific Ocean (Marincovich et al., 1990). The abrupt transition tothe present day Holocene or interglacial mode conditions�10 K years ago marks an almost complete shift in variables anddrivers governing the Arctic system. During the Holocene, onshorter time scales of 1000–100 years, significant changes in highlatitude climate are now becoming apparent in temperature (e.g.the medieval warm period and little ice age) and perhaps moisturedelivery patterns (e.g., Overpeck et al., 1997). The pace of changeover the past three decades has been even more impressive, withsummer ice extent in 2012 being 54% less than the 1980 bench-mark and thickness diminishing by half (Cavalieri and Parkinson,2012; Barber et al., 2015 and references therein). Associated withthe rapid retreat and thinning of sea ice the water column haswarmed at depths exceeding 800 m owing to warmer watersentering from the Atlantic and Pacific Oceans (Polyakov et al.,2013; Table 2).

The truly unique aspect of such change for the Arctic basins isthat now, for the first time in the observational record, sea ice isretreating on an annual basis past the shelf break, and thus expos-ing basin waters directly to sunlight and wind forcing. Our roughcalculation suggests that the summer 2012 sea ice retreat leftapproximately 40% of the basin area exposed. One consequenceis that upwelling favorable winds can now directly and efficientlydrive shelf-break upwelling, and draw nutrients from subsurfacebasin waters onto the shelf (Section 6.3) (Carmack and Chapman,2003). However, the same upwelling favorable winds will also cre-ate onshore pressure gradients over the slope and basin, which willact to disrupt the flow of waters in the ACBC (see Section 6.3).Within the basin interior remaining ice is thinner and less com-pact, and thus more responsive to wind stress (forcing and mixing)(Kwok et al., 2013).

Within the EB, and later in the AB, the first observed changes inwater column temperature were advection related (Quadfaselet al., 1991; Carmack et al., 1995). Since then both observationalevidence and modeling results suggest that temperature changesin AW are related to pulses in the AW inflows (Polyakov et al.,2005; Karcher et al., 2011). The first signal of far-field effects onCanada Basin waters, as delivered by the ACBC, was the increased

Table 2Observed and modeled changes in the Amerasian and Eurasian Basins in the past few decades.

Layer Amerasian basin (mostly Pacific influenced part of Canada basin) Eurasian basin

Sea ice – Remaining MYI limited to north of Canadian Archipelago(Stroeve et al., 2012; Barber et al., 2015)

– Remaining MYI limited to north of Greenland (Barber et al.,2015)

– Ice algal production from shelves transported to basins, wherefast vertical flux supplied food to benthic fauna (Boetius et al.,2013)

Surface mixed layer – Decreasing salinity, increased stratification, deepening nutri-cline (McLaughlin and Carmack, 2010; Jackson et al., 2010)

– Observed and modeled high variability in freshwater distribu-tions and pathways (Newton et al., 2008, Polyakov et al. 2008,Morison et al., 2012)

– Very low nutrient concentration in Beaufort Gyre suggests no/little increase in primary production is possible (cf. Codispotiet al., 2013)

– Small phytoplankton thrive (Li et al., 2009)– Decrease for small zooplankton dependent on authochtonous

production, increase/no change in large zooplankton benefit-ing from allochtonous production (Hunt et al., 2014)?

– Aragonite-saturation decreasing (Yamamoto-Kawai et al.,2009)

– Variability in freshwater content (Timmermans et al., 2011)– Nutrient concentrations not limiting, perhaps increase in pri-

mary production possible (Codispoti et al., 2013)– No change in primary production between 1995 and 2007

(Wassmann et al., 2010)– Increased frequency of fall blooms (ACIA, 2005; Ardyna et al.,

2014)– (modeled) increase in secondary production across both

basins, slightly higher in Eurasian Basin (Slagstad et al.,2011, their Fig. 9)

Halo-cline – Enhanced shelf-break upwelling (Carmack and Chapman,2003; Tremblay et al., 2011; Pickart et al., 2013)

– Interannual shifts in location of the Atlantic-Pacific haloclinefront (McLaughlin et al., 1996; Karcher et al., 2012)

– Possible warming owing to increased vertical heat flux fromAW (Polyakov et al., 2012)

Atlantic water – Warm pulse penetration into the basin (Shimada et al., 2004)and temperature increase between 0 and <1 �C (McLaughlinet al., 2009; Polyakov et al., 2013)

– More intense anticyclonic Beaufort Gyre after 2004 constrain-ing the cyclonic ACBC that transports AW (Proshutinsky et al.,2009; Karcher et al., 2012)

– Temperature increase by �1 �C with change in inflow volume(Polyakov et al., 2013)

Arctic Basins Deep water – Geothermal warming (0.004 �C per decade; Carmack et al.,2012), no time series to document biological change

– Warming more than expected from geothermal heating(Rudels et al., 2013)

– Decreased megafauna densities and sediment microbial bio-mass at HAUSGARTEN for 2002–2007 (Bergmann et al., 2011)


ventilation in the 1990s of Atlantic-origin BSB waters as shown byfreshening and increased levels of chlorofluorocarbon and oxygenat depths of 600–1600 m (McLaughlin et al., 2002). WarmerAtlantic-origin FSB water, first observed in the Nansen Basin duringthe early 1990s, extended to the Northwind Ridge along the west-ern reaches of the Canada Basin in 2002 (Shimada et al., 2004) andsubsequently spread across most of the southern basin interior by2007 (McLaughlin et al., 2009). These newer and warmer FSBwaters were also fresher and more ventilated than existing basinwaters. A second pulse was observed to begin in Fram Strait in1999 (Schauer et al., 2004) and then appeared in the eastern EBin 2004; Polyakov et al. (2005) used this warming pulse to calcu-late an advection speed of along-slope warming of 1.5 cm s�1.The warming pulse peaked in the Arctic basin interior in 2007–2008 with maximum temperature anomalies of up to 1 �C, and2008–2010 observations suggest that the AO interior was in tran-sition toward a cooler state in the 1990s, but not well-correlatedwith increases in the temperature of Bering Sea source waters(Polyakov et al., 2011). Shimada et al. (2006) noted that recentincreases in Pacific summer water temperatures arewell-correlated with the onset of sea-ice reduction that began inthe late 1990s. They suggest that reduced internal ice stress inautumn allows a more efficient coupling of anti-cyclonic windforcing to ice and the upper ocean, which subsequently redirectsPacific summer water into the central basin at sufficiently shallowdepths to affect winter heat loss to the atmosphere and ice.

In concert with increasing focus on tracking sea ice changes inthe Arctic over the past few decades, the role of the AO halocline,its possible variability over time and its different structure in theAB and EB have received increasing attention. Increased studyeffort is primarily due to the climate-relevant function of the halo-cline as the insulating layer between the cold polar mixed layer

above and the comparatively warm Atlantic layer underneath(e.g. Shimada et al., 2005). In this sandwiched position, the halo-cline controls the extent to which the underlying and warmerbasin waters affect sea ice volume and phenology. A strengtheningin the stratification of the halocline in the Canada and MakarovBasins was observed through a decade-long series of CTD databetween 1997 and 2008, whereas the halocline was apparentlyrather stable in the Amundsen Basin over that same time period(cf. Bourgain and Gascard, 2011). Temporal changes within thesummer halocline in the Canada Basin are related to developmentof a near-surface temperature maximum within the halocline andshoaling of the summer halocline (Jackson et al., 2010). Theseauthors argue that heat contained within the shallownear-surface temperature maximum and released during thefreeze-up season could delay freeze-up or reduce ice thickness.

Light regime and nutrient supply critically influence theamount of primary production that can take place in the basins.How will nutrient concentrations play out in the future ArcticBasins? In their modeling study, Zhang et al. (2010) predict anincrease in nutrient availability to the generalized upper AObecause of enhanced air-sea momentum transfer due to reducedsea ice cover. Although these authors do not specify exactly whereor how widespread the nutrient replenishment would occur, nutri-ent enhancement will likely occur primarily along the shelf breakrelated to upwelling, rather than in the basins per se (Carmackand Chapman, 2003; Tremblay et al., 2011; Table 2). Furthermore,horizontal advection of nutrient rich PW and AW may contributethe nutrients for up to 20% of the total AO primary production(Popova et al., 2013). For the basins, model predictions of enhancednutrient availability do not match the observed reduced nitrateconcentrations of the Canada Basin time series of the upper ocean(Li et al., 2009) or recent observations of extremely low nutrient


concentrations in the same area (Codispoti et al., 2013; Matrai andApollonio, 2013).

More numerous than observations and predictions of changingnutrient concentrations are observations and predictions about thetrends in primary production, although some failed to separate theArctic basins from the shelves and some seem contradictory. Sev-eral authors agree that primary production in the AO as a wholewill increase – or has increased - under current climate change sce-narios, but with significant regional variability and including areasof no increase or decrease. Exactly where and by how much pri-mary production may change has been estimated based on models,in situ observations and satellite imagery (e.g. Arrigo et al., 2008;Pabi et al., 2008; Zhang et al., 2010; Slagstad et al., 2011; Ji et al.,2013; Vancoppenolle et al., 2013), but is debated in light of thedependence on and variability in light and nutrient availabilityand re-supply, and extent and decline in sea ice cover (Tremblayet al., 2009; Tremblay and Gagnon, 2009; Lee et al., 2010; Zhanget al., 2010; Brown and Arrigo, 2012; Codispoti et al., 2013). Theprimary reasons brought forward to explain increases in primaryproduction both on the shelves and in the basins include increasingduration of the open water period and a growing spatial extent ofopen water, resulting in longer and greater exposure to solar radi-ation with a resulting extended phytoplankton growing season(e.g. Arrigo et al., 2008; Zhang et al., 2010; Slagstad et al., 2011;Ardyna et al., 2014). Estimates based on remotely sensed chloro-phyll, sea ice cover and sea surface temperatures suggest that halfof the increase in primary production between 1998 and 2006occurred in ‘‘pelagic waters’’ (Pabi et al., 2008), but no estimateshelf versus basins waters was given. Similarly, modeled primaryproduction in the upper 100 m for the period of 1998–2007increased mostly ‘‘in the permanently and seasonally-coveredareas with decreasing ice thickness and increasing summer meltback in recent years’’ [versus the open waters of the GIN Seasand southern Barents Sea] (Zhang et al., 2010), but the proportionsof increase over the basins versus the shelves were, again, not spec-ified. A different model estimated a basin-specific 4-fold increaseof gross primary production assuming an air temperature increaseof up to 8 �C (Slagstad et al., 2011). An inter-comparison betweenfive different biogeochemical models for the Arctic summarizedthat physical and chemical boundary conditions differed substan-tially between models, mainly in terms of nutrient concentrationsand winter mixing (Popova et al., 2012), and these authors cautionthat predictive capabilities of models are reduced as long as thesebasic boundary conditions are inconsistent.

Comparisons of modeled phenology of primary and sec-ondary (micro-, meso-, predatory zooplankton) production atdifferent locations along the SHEBA drift (Ashjian et al., 2003)showed that spatial and temporal variability in primary productioncan be expected to be largest in areas with seasonal icecover (Zhang et al., 2010). Given that seasonal ice cover is pre-dicted for the entire basin before the end of the century (Wangand Overland, 2009), substantial increases in primary productionand changes in phenology might be inferred from the Zhanget al. (2010) results. Another, basin-specific forecast, however,actually suggests low or no increase in primary production forthe AB, particularly the Canada Basin, and northern Beaufort Seabecause of particularly strong stratification and very low nutrientconcentrations (Bluhm and Gradinger, 2008; Codispoti et al.,2013). Although nutrient (nitrate) concentrations do not appearto be limiting in the EB (Codispoti et al., 2013), modeled primaryproduction for the period 1995–2007 did not increase in the Euro-pean Basin sector (Wassmann et al., 2010). Regardless of the trajec-tory of the quantity of primary production, increased lightintensity below the ice and in open water will favor a longer veg-etative season (Arrigo et al., 2012; Ji et al., 2013; Ardyna et al.,2014).

The anticipated change in ice algal production has received lessattention than phytoplankton production. In the 1990s, ice algaecontributed a larger proportion to total algal biomass in the ArcticBasins (up to 50%) than over the shelves (Gosselin et al., 1997). Thehigh ratio in the Arctic Basins is mainly driven by the very lowpelagic primary production in the oligotrophic Arctic (Section 4.2).The relatively high pelagic primary production extant on the Arcticshelves may explain why model-estimated annual Arctic phyto-plankton production in the upper ocean from 1998–2006 wasfound to be �20 times higher than that of ice algal production(Jin et al., 2012), and thus a distinction must be made betweenshelf and basin ice algal production.

Algal production associated with sea ice is by many assumed todecrease in the future based on the overall decrease in sea ice cover(e.g. Zhang et al., 2010). It is worth asking, however, if the antici-pated decline could be counterbalanced, in the near future at least,by the switch from multi-year to first-year sea ice in the ArcticBasins (Barber et al., 2015), which has several effects. First, thinnerice allows more light to penetrate the ice resulting in higher icealgal production than in thicker ice, provided nutrients are avail-able for growth (Gradinger, 2009), and snow depth does notincrease. Second, nutrient exchange across the ice-water interfaceis more efficient in warmer, saltier and thinner first-year ice thanin the multi-year ice of the past (Deal et al., 2011). Third, surfacemelt ponds and melt holes producing a patchwork of small (belowsatellite detection) but productive hot spots have been suggestedto become more common in the central Arctic in the future (Leeet al., 2011). And last, massive amounts of ice algae presumablygrown on the shelves have recently been observed over the centralbasin where they contributed to algal biomass (Boetius et al.,2013). In the end, the balance of sea ice and ice algal bloom phenol-ogy with regional nutrient distribution patterns will determine thetrends in absolute and relative ice algal production (Leu et al.,2015).

What do the outlined physical and biological changes mean forthe pelagic biota in the Arctic basins, vertical carbon flux, and theultimate carbon recipients, the Arctic deep-sea benthos? For zoo-plankton, one can reasonably anticipate that the production ofautochthonous zooplankton will closely be tied to future primaryproduction levels in the Arctic Basins, and allochthonous zooplank-ton biomass will continue to be driven by the inflow of Atlanticand Pacific Waters (Wassmann et al., 2015). Model estimates sug-gest increasing secondary production (based on Calanus glacialis) inthe basins from an average zero or negative rate up to 1.5 g C m�2 -y�1 under their extreme scenario of 8 �C air temperature increase(Slagstad et al., 2011). Given the uncertainty of the trend and tra-jectory in primary production for the basins, however, (see dis-agreement between models on primary forcing factors in Popovaet al., 2012) and the resulting particulate organic matter poolsavailable for food webs, this trend seems uncertain. Future statesfor ice-associated invertebrate fauna are equally uncertain, butare likely tied to sea ice extent and ice algal production (Bluhmet al., 2010). Survival strategies for low ice conditions may exists;Berge et al. (2012) hypothesize that some ice fauna may conductvertical migrations to depths where deep northward currentswould take them back into sea ice-covered areas where theyascend again and re-colonize the underside of drifting pack icewhen present. These authors suggest this mechanism wouldreduce loss of ice fauna from the Arctic.

Predictions on vertical flux and resulting benthic standingstocks currently remain largely speculative. Two opposing scenar-ios are: Increasing primary production linked to decreasing sea icecover, rising surface water temperature and enhanced zooplanktonproduction (cf Pabi et al., 2008; Zhang et al., 2010; Slagstad et al.,2011) could result in increasing vertical export of particulateorganic matter to the sea floor. This scenario finds support in a


2-year sediment trap deployment over the Laptev Sea marginwhere POC export flux increased in response to the extreme sum-mer sea ice melt in 2007 (Lalande et al., 2009). The opposing sce-nario could be that vertical flux in the basins may be hamperedby increasing vertical stratification (at least in the Canada Basin;Carmack et al., 2006; Jackson et al., 2010), low and perhaps declin-ing nutrient concentrations in the upper layer of the Canada Basin(Li et al., 2009; Codispoti et al., 2013), and the co-variance of lightweight, small algae with strong stratification (Li et al., 2009; Leeet al., 2012). Presumably those small algae would have limitedpenetration capability in a strengthening pycnocline compared tothat of heavy, large, chain-forming diatoms as described byBoetius et al. (2013). Along the same lines, a shift from large dia-toms to smaller coccolithoporid cells occurred at the deep-seaobservatory HAUSGARTEN in Fram Strait over the last decade inwarm years, and vertical carbon flux of biogenic particulate silicawas lower in warm than in cold years (Lalande et al., 2013).Changes in water column processes in some fashion perpetuatedto epibenthic megafauna densities that declined at HAUSGARTENbetween 2002 and 2007 with concurrently declining sedimentorganic content, sediment microbial biomass and increasing bot-tom water temperatures (Bergmann et al., 2011). Temporal trendson vertical flux or of deep benthic communities in the AB are, toour knowledge, not available.

8. Outlook

Throughout much of the 20th century the AO – and its centraldeep basins in particular – was viewed as a small, remote, slowlychanging and relatively unimportant part of the global system. Aperceived need to catch-up for lost time now prevails, requiringthat we pay more attention to the pan-Arctic perspective and tothe importance of advection and physical/biological connectivityand their joint roles for the basins. This paper summarizes, froman interdisciplinary point of view, our current perception of howthe Arctic Basins are set-up and operating, and how basin oceanog-raphy and biology have recently (e.g. the last few decades) changed(Table 2). The AO is subject to large amplitude multi-decadal vari-ability and long-term trends (Polyakov et al., 2013), thus challeng-ing interpretations of observed changes to climate drivers. Theseauthors state, however, that the exceptional magnitude of recenthigh-latitude changes, both oceanic and atmospheric, implies anirreversible shift of the AO to a new climate state (see alsoJeffries et al., 2013; Wood et al., 2013). What does this ‘new’ statehold for biota? A review of the climate change ‘footprints’ in arcticecosystems by Wassmann et al. (2011) gives evidence that all com-ponents of the high-latitude marine ecosystem are being impactedby global change, though most existing reports considered largemammals and birds only, and mainly in shelf areas.

What, then, do we know? Concrete findings from decade-longtime-series span from declines in sea ice extent and thickness(Kwok et al., 2009; Stroeve et al., 2012, Barber et al., 2015) toincreasing river discharges (McClelland et al., 2006), with conse-quences that include the appearance of aragonite-undersaturated(low pH) waters in the Canada Basin this past decade(Yamamoto-Kawai et al., 2008, 2011; Table 2). Other changesinclude warming of both Atlantic (Polyakov et al., 2011) and Pacific(Shimada et al., 2006) inflow waters and increased stratificationresulting from both ice melt and increased river inflow (Jacksonet al., 2010). Primary production appears to have increased overthe shelves (Pabi et al., 2008; Zhang et al., 2010) but perhaps notover the basins because of – at least regionally – nutrient limitation(Tremblay et al., 2011; Codispoti et al., 2013; Matrai andAppolonio, 2013) along with changes in algal community sizestructure with prevalence of small taxa in nutrient-poor basin

areas or times (Li et al., 2009; Lee et al., 2012). Knowledge oflong-term patterns of vertical flux and biological observationsabove the primary producer level, however, is rudimentarybecause of very short time series (Bauerfeind et al., 2009;Bergmann et al., 2011). The interconnections between physical,chemical and (lower trophic) biological changes are slowly begin-ning to be incorporated into pan-Arctic models and clearly docu-ment that such connections exist, although they yet need to betied to higher tropic levels (Wassmann et al., 2010 and 2015,Zhang et al., 2010; Slagstad et al., 2011).

One basic question remains elusive: will new (export) primaryproduction increase or decrease under conditions of a temporallyand spatially reduced ice cover and what role will the basins play?And will the Arctic food web of the future provide more or lessenergy to the higher trophic levels, with potential implications tothe development of commercial fisheries? The response will likelydepend on region, process and scale. On the panarctic shelves,which we include here for comparative purposes the annualremoval of ice cover beyond the shelf break will enhance windand ice-forced shelf-break upwelling; the resulting increases inboth nutrient fluxes and solar radiation should thus increase newproduction (Tremblay et al., 2011; Williams and Carmack, 2015and Section 6.3); however, the stronger winds required to drivesuch upwelling are more common in fall during diminishing lightconditions. This may result in a fall bloom setting (cf. ACIA, 2004,Fig. 9.13, Ardyna et al., 2014) or result in delayed uptake untilthe following spring. In the deep central basins the increased addi-tion of freshwater from ice melt will increase surface layer stratifi-cation, particularly in the Pacific Arctic sector (cf. Jackson et al.,2010). In the Beaufort Gyre, increased coupling of wind stress tosurface waters under reduced ice cover has increased Ekman con-vergence of surface waters, and thus further increased stratifica-tion (Proshutinsky et al., 2009; McLaughlin and Carmack, 2010).In this case the resulting decrease in vertical nutrient flux to theeuphotic zone should thus decrease new production, but perhapsincrease the relative contribution of ice algae.

The current distribution of faunal biomass (zooplankton, ben-thos, fish) suggests the energy transfer from primary productioninto secondary production in the basins is (1) generally concen-trated at the shelf break, i.e. along the basin perimeter, withdecreasing trends associated with increasing depth, and (2) peaksin areas receiving influx of Atlantic and Pacific -source nutrients,new production and both allochthonous and autochthonous fauna.Abutment of water masses along the slope (e.g. the Atlantic layer incontact with the bottom) may provide preferential habitat forsome species during all or some portion of their life cycle, andmechanisms of pelagic/benthic coupling (Wassmann, 1998;Grebmeier, 2012) are thus of paramount importance. For example,concentrations of Arctic/polar cod (Boreogadus saida), appear topeak in areas where the (upper portion of) the AW meets the seafloor (Logerwell et al., 2011; Parker-Stetter et al., 2011; Crawfordet al., 2011). Assuming this distribution pattern is found to holdtrue around the basin perimeter, one needs to be concerned aboutthe potential for these fish providing an easy and concentrated tar-get for future fisheries in the Arctic Basins, thus requiring newmanagement policies in multiple nations (Christiansen et al.,2014). With various fish species changing their core concentrationsnorthward in the northern hemisphere (Mueter and Litzow, 2008;Perry et al., 2005), fisheries fleets are certain to follow suit; on theBarents Sea shelf a B. saida fishery already goes back to the 1960s(summarized by Hop and Gjøsæter, 2013). Easy access of a fleet toharvestable grounds would have cascading effects throughout thefood web, reaching to vertebrate predators given that B. saida hasbeen a stable and lipid-rich staple (c.f. Hop and Gjøsæter, 2013)for those subsistence-harvested, poster-child Arctic marinemammals.


In summary, for the first time in recent history, a new deepocean appears to be opening (cf. Kinnard et al., 2011) and it is likelythat within a few decades or less the Arctic will see mostly ice-freesummers extending fully across its basins. The rapidly changingArctic, leading the globe through poorly understood climate ampli-fication processes, offers much in return to allow in-depth study ofsocial and ecological processes that will be impacted in a warmerworld (ARR, 2013). Prediction of future states is challenging.Rampal et al. (2011) showed that models underestimate theobserved thinning trend by almost a factor 4 on average, in away even more spectacular than they do for the sea ice extentdecline and this critical underestimation of the thinning trend can-not be explained entirely by an underestimation of the decline ofthe perennial, thicker, ice-covered portion of the Arctic. Itsmediterranean configuration, with connections to the subarcticAtlantic and Pacific, provides an ideal setting in which to addressecology in an advective system (Wassmann et al., 2015). Its deepbasins, in particular, are experiencing conditions that have notoccurred in hundreds if not thousands if not hundreds of thousandyears. As a leading indicator of change it is, in fact, a window to thefuture, providing an opportunity to unite science and policy inways that can be tested in the high-latitudes and the later appliedwith more confidence to the lower latitudes (cf. Carmack et al.,2012). Although uniting science and policy is challenging giventhe different agendas and interests between policy makers and sci-entists, it was in part the political information needs that moti-vated numerous basin-focused research investigations in the pastdecade such as mapping activities that were driven by the desireto clarify national boundaries. Building on the achievements ofIPY and the International Conference on Arctic Research PlanningII-outcomes (ICARP II Science Plan 4, 2005) and developing ICARPIII, the Arctic research community has shown and will continueto show pathways forward on how international collaboration ismore productive than focusing on national boundaries. This timearound, the central AO will play an increasing larger role both inthe research and political arenas.

Acknowledgments

This article resulted from the 3rd pan-Arctic Symposium enti-tled ‘‘Overarching perspectives of contemporary and future ecosys-tems in the AO’’ held in Motuvun, Croatia, in October of 2012. Wethank the workshop organizers P. F. Wassmann and M. Reigstad forcreating an inspiring atmosphere for fruitful discussion, and forsupporting the authors’ travel. We appreciate P.F. Wassmann’s per-sistent encouragement as our guest editor. T. Kikuchi and collabo-rators are thanked for letting us use their data in Fig. 7. Themanuscript benefited from discussion with B. Björk, R. Crawford,R. Gradinger, I. Polyakov, R. Rember, and K. Wood. We gratefullyacknowledge the preparation of graphics by Falk Huettmann, Patri-cia Kimber and Liusen Xie. Mette Kaufman and Flora Grabowskahelped with the reference list. KK was partially supported by theRussian Foundation for Basic Research under Grant 13-04-00551.This review has benefited from BB’s work funded by NOAA’s Officeof Ocean Exploration (grant #NA 16RP2627), The Bureau of OceanEnergy Management (Agreement #M12AC00011), and the Alfred.P. Sloan Foundation’s Arctic Ocean Diversity Census of Marine Lifeproject.

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A Tale of Two Basins an Integrated Physical and Biological Perspective of the Deep Arctic Ocean 2015...

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