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Surface freshening in the Arctic Ocean's Eurasian Basin: An apparent consequence of recent change in the wind-driven circulation

2011; American Geophysical Union; Volume: 116; Linguagem: Inglês

10.1029/2011jc006975

2156-2202

Mary‐Louise Timmermans, Andrey Proshutinsky, Richard Krishfield, Donald K. Perovich, J. Richter‐Menge, Timothy P. Stanton, John M. Toole,

Cryospheric studies and observations

Journal of Geophysical Research: OceansVolume 116, Issue C8 Free Access Surface freshening in the Arctic Ocean's Eurasian Basin: An apparent consequence of recent change in the wind-driven circulation M.-L. Timmermans, M.-L. Timmermans [email protected] Department of Geology and Geophysics, Yale University, New Haven, Connecticut, USASearch for more papers by this authorA. Proshutinsky, A. Proshutinsky Department of Physical Oceanography, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USASearch for more papers by this authorR. A. Krishfield, R. A. Krishfield Department of Physical Oceanography, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USASearch for more papers by this authorD. K. Perovich, D. K. Perovich Cold Regions Research and Engineering Laboratory, Hanover, New Hampshire, USASearch for more papers by this authorJ. A. Richter-Menge, J. A. Richter-Menge Cold Regions Research and Engineering Laboratory, Hanover, New Hampshire, USASearch for more papers by this authorT. P. Stanton, T. P. Stanton Department of Oceanography, Naval Postgraduate School, Monterey, California, USASearch for more papers by this authorJ. M. Toole, J. M. Toole Department of Physical Oceanography, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USASearch for more papers by this author M.-L. Timmermans, M.-L. Timmermans [email protected] Department of Geology and Geophysics, Yale University, New Haven, Connecticut, USASearch for more papers by this authorA. Proshutinsky, A. Proshutinsky Department of Physical Oceanography, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USASearch for more papers by this authorR. A. Krishfield, R. A. Krishfield Department of Physical Oceanography, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USASearch for more papers by this authorD. K. Perovich, D. K. Perovich Cold Regions Research and Engineering Laboratory, Hanover, New Hampshire, USASearch for more papers by this authorJ. A. Richter-Menge, J. A. Richter-Menge Cold Regions Research and Engineering Laboratory, Hanover, New Hampshire, USASearch for more papers by this authorT. P. Stanton, T. P. Stanton Department of Oceanography, Naval Postgraduate School, Monterey, California, USASearch for more papers by this authorJ. M. Toole, J. M. Toole Department of Physical Oceanography, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USASearch for more papers by this author First published: 23 July 2011 https://doi.org/10.1029/2011JC006975Citations: 77AboutSectionsPDF ToolsRequest permissionExport citationAdd to favoritesTrack citation ShareShare Give accessShare full text accessShare full-text accessPlease review our Terms and Conditions of Use and check box below to share full-text version of article.I have read and accept the Wiley Online Library Terms and Conditions of UseShareable LinkUse the link below to share a full-text version of this article with your friends and colleagues. Learn more.Copy URL Share a linkShare onFacebookTwitterLinkedInRedditWechat Abstract [1] Data collected by an autonomous ice-based observatory that drifted into the Eurasian Basin between April and November 2010 indicate that the upper ocean was appreciably fresher than in 2007 and 2008. Sea ice and snowmelt over the course of the 2010 drift amounted to an input of less than 0.5 m of liquid freshwater to the ocean (comparable to the freshening by melting estimated for those previous years), while the observed change in upper-ocean salinity over the melt period implies a freshwater gain of about 0.7 m. Results of a wind-driven ocean model corroborate the observations of freshening and suggest that unusually fresh surface waters observed in parts of the Eurasian Basin in 2010 may have been due to the spreading of anomalously fresh water previously residing in the Beaufort Gyre. This flux is likely associated with a 2009 shift in the large-scale atmospheric circulation to a significant reduction in strength of the anticyclonic Beaufort Gyre and the Transpolar Drift Stream. Key Points Eurasian Basin upper ocean was appreciably fresher in 2010 than in 2007–2008 Change in salinity due to local melt cannot account for the anomaly Observed fresh water attributed to a 2009 shift in the atmospheric circulation 1. Introduction [2] Fresh water maintains the strong near-surface Arctic Ocean stratification that inhibits transfer of deep ocean heat to the surface, and consequently has a major impact on sea ice cover, and the entire Arctic climate system. The spatial distribution of liquid fresh water is controlled by: sea ice growth and decay; advection of fresh river water and relatively low salinity Pacific water from the Arctic marginal seas; precipitation; and wind forcing, which affects sea ice drift and water-mass pathways and modulates the depth of the halocline. These processes contribute to the time-varying liquid freshwater content to varying degrees in the different Arctic Ocean regions and over seasonal to decadal time scales [e.g., Steele and Ermold, 2004; Serreze et al., 2006; Polyakov et al., 2008; Newton et al., 2008; Dmitrenko et al., 2008; Rabe et al., 2010]. In this paper, we examine recently observed substantial changes in the Eurasian Basin mixed-layer salinity in the context of seasonal processes and the large-scale atmospheric circulation. The mechanisms of the coupled Arctic ocean-sea ice-atmosphere system examined here offer a foundation for model validation and improvement, contributing to the Arctic Ocean Model Intercomparison Project (AOMIP) community. [3] Changes in the freshwater distribution of the upper ocean have been shown to be linked to major shifts in Arctic wind forcing regimes that occur on time scales of several years to decades [e.g., Morison et al., 1998; Steele and Boyd, 1998; Anderson et al., 2004; Morison et al., 2006; Alkire et al., 2007; Newton et al., 2008]. Upper-ocean salinity changes are particularly large in the central Arctic Ocean in the vicinity of the front between less saline water masses of Pacific character and more saline water masses of Atlantic character [e.g., Morison et al., 1998; McLaughlin et al., 2004]. Rapid changes (on interannual, and possibly seasonal, time scales) in upper-ocean freshwater content in parts of the central Arctic Ocean take place as a result of shifts in this frontal zone. For example, Morison et al. [1998] showed that in the 1990s the front shifted from roughly over the Lomonosov Ridge (based on pre-1990s climatology) to roughly over the Alpha and Mendeleyev Ridges, and that this change in the ice and upper-ocean circulation could be attributed to a shift in atmospheric circulation. At the same time, the cold Arctic halocline disappeared from the Eurasian Basin, while mixed layers became more saline there [Steele and Boyd, 1998]. In the 2000s, upper-ocean salinity and temperature in the central Arctic Ocean returned to approximately pre-1990s conditions [Morison et al., 2006] and the Eurasian Basin saw a partial return of the cold halocline [Boyd et al., 2002]. [4] Similarly, the atmospheric circulation effectively regulates the strength and position of the Beaufort Gyre circulation system, centered in the Canada Basin. Owing to prevailing anticyclonic (clockwise) winds generated by the Arctic high, the Beaufort Gyre accumulates fresh water from sea ice melt, Pacific Water inflows through Bering Strait, river runoff and atmospheric precipitation in the upper ocean primarily by Ekman convergence [e.g., Proshutinsky et al., 2002]. There is a direct correlation between the size and intensity of the Arctic high, position of the frontal zone discussed above, freshwater content in the Beaufort Gyre region and changes in upper-ocean salinity in the Eurasian basin. The large-scale atmospheric winds redistribute water masses such that an increase of fresh water in the Beaufort Gyre results in a deficit of fresh water (i.e., an upper-ocean salinity increase) in the Eurasian sector of the Arctic Ocean and vice versa. Aagaard and Carmack [1989] showed that a deeper halocline and fresher surface waters in the Canadian Basin result in about 4 times more liquid fresh water (relative to a salinity of 34.8) than in the Eurasian Basin. Recent years have seen an unprecedented increase of fresh water accumulated in the Beaufort Gyre: specifically, between 2003 and 2009, the freshwater content of the Beaufort Gyre increased by more than 30% [Proshutinsky et al., 2009]. Locally, McPhee et al. [2009] show that in parts of the southeast Canada Basin, freshwater content in 2008 was as much as 60% above the climatological values. Note that McPhee et al. [2009] found a significant reduction in freshwater content in the Makarov and Amundsen basins. [5] A conceptual model put forward by Proshutinsky et al. [2002] asserts that the Beaufort Gyre accumulates a considerable volume of fresh water under large-scale anticyclonic wind forcing (high surface atmospheric pressure), and releases this fresh water when the forcing weakens or changes sense to cyclonic. In this scenario, anticyclonic winds tend to accumulate fresh water in the Beaufort Gyre by Ekman convergence while the cyclonic circulation regime is associated with Ekman divergence and a reduction of fresh water in the Gyre [see also Hunkins and Whitehead, 1992]. In addition to this dynamic forcing, thermodynamic forcing (sea ice growth and decay) also regulates Beaufort Gyre liquid fresh water. Proshutinsky et al. [2009] showed that the mean seasonal cycle of ocean fresh water in the Beaufort Gyre has two peaks: one in summer when the sea ice thickness reaches its minimum (maximum ice melt), and the other in late fall–early winter when the wind stress curl is strongest (maximum Ekman pumping) while the salt input from growing sea ice has not yet reached its maximum. Between 1997 through 2008, the Arctic was characterized by a strong anticyclonic circulation. The unusually long 12 year anticyclonic conditions were likely a major factor in the substantial increase in Beaufort Gyre fresh water over this time [Proshutinsky et al., 2009; McPhee et al., 2009]. [6] Freshwater content variability is further impacted by changes in freshwater sources [e.g., Peterson et al., 2002]. For example, Eurasian river runoff to the Arctic was 25% larger in 2007 than the mean river influx between 1936 and 2006 [Shiklomanov and Lammers, 2009]; these sizable variations are superimposed on local freshwater changes arising from changes in the large-scale circulation. At the same time, river discharge is influenced by the large-scale atmospheric circulation. Proshutinsky et al. [2000] show that during anticyclonic wind regimes, atmospheric cyclone trajectories are shifted toward Siberia, bringing increased precipitation that feeds Siberian rivers. This is in contrast to cyclonic wind regimes when cyclones propagate preferentially to the Central Arctic, and Siberian river discharge decreases. The most intense annually averaged anticyclonic circulation in the last 60 years was seen in 2007 [Proshutinsky and Johnson, 2010], consistent with the observed maximum river discharge [Shiklomanov, 2010; Shiklomanov and Lammers, 2009]. [7] In this study, we examine the consequences of a significant change in 2009 in the wind-driven circulation (on a background of seasonal melt) to the surface ocean freshwater content in the Eurasian Basin. We assess the evolution and spatial distribution of fresh water in the ocean mixed layer by analyzing ocean and sea ice measurements from an ice-based observatory (IBO) that drifted in the Eurasian Basin in 2010, together with hydrographic data collected in the same region in previous years. We interpret our observational results in the context of the large-scale circulation by employing a wind-forced Arctic Ocean model with frictional coupling between the ocean and sea ice [see Proshutinsky and Johnson, 1997]. The model is used to examine the transport of fresh surface waters from the Beaufort Gyre and Siberian shelf regions to the Eurasian Basin and Fram Strait, and to provide support to our hypothesis on the origin of observed fresh water in the Eurasian Basin. In section 2, we investigate seasonal melt and the freshwater content of surface waters measured by an IBO that was deployed on a 1.7 m thick ice floe in the Transpolar Drift on 19 April 2010 at 88°39.4′N, 145°35.7′E in conjunction with the North Pole Environmental Observatory (NPEO) program. The IBO included a Woods Hole Oceanographic Institution ice-tethered profiler (ITP), a U.S. Army Cold Regions Research and Engineering Laboratory ice mass balance buoy (IMB), a Naval Postgraduate School Arctic Ocean flux buoy (AOFB), and a National Oceanic and Atmospheric Administration (NOAA)–Pacific Marine Environmental Laboratory (PMEL) Web camera. The combined ITP and IMB instrumentation allows us to estimate the surface ocean freshening due to ice and snowmelt over the course of the IBO drift (Figure 1) and to relate this to concurrent upper-ocean temperature and salinity measurements. In addition to estimating bounds on freshening due to local melt, we compare data from previous years to show that the entire IBO drift region was significantly fresher in 2010 compared to 2007 and 2008 even though seasonal melt was comparable in all 3 years. In section 3, we employ the two-dimensional model to explore the wind-driven sea ice and upper-ocean Arctic circulation in the context of the observations. In section 4, results are summarized and discussed in perspective with past upper-ocean circulation changes in the region. Figure 1Open in figure viewerPowerPoint Drift track of the ice-based observatory (IBO) between 19 April 2010 and 10 November 2010, indicating ice-tethered profiler (ITP) salinity at 10 m. The 1000, 2500, and 3000 m isobaths have been plotted using the International Bathymetric Chart of the Arctic Ocean (IBCAO) grid. Images are from the NOAA Pacific Marine Environmental Laboratory (PMEL) Web camera installed on the same floe. Black pointers indicate the IBO location at the start and end of the melt period. 2. Upper-Ocean Freshening IBO Measurements [8] The Ice-Mass Balance Buoy consists of a thermistor string extending from the surface through the snow and ice into the upper ocean (typically between 200 and 380 cm below the top surface of the ice), acoustic sounders above the ice-snow surface and below the ice bottom, GPS, barometer, air temperature sensor, and an Iridium transmitter for data recovery [Richter-Menge et al., 2006] (http://imb.crrel.usace.army.mil/). The ITP system consists of a surface buoy deployed in the ice floe, and an automated profiling CTD instrument that is mounted to a tether that extends to about 750 m below [Krishfield et al., 2008] (http://www.whoi.edu/page.do?pid=20756). The profiling CTD cycles vertically along the tether from about 7 m depth to about 750 m; the systems examined here returned four one-way profiles of temperature and salinity (at about 25 cm vertical resolution) per day via Iridium satellite (data are available on the ITP Web site at http://www.whoi.edu/itp/data/). The Autonomous Ocean Flux Buoy consists of a surface buoy that sits on the ice and an instrument frame suspended into the upper ocean. The instrument frame includes an acoustic travel time current meter (Falmouth Scientific Inc., FSI) located nominally 4 m below the ice that measures three component velocities with mm s−1 noise levels. A submillidegree resolution thermistor string with sensors spaced 0.4 m apart extended up from the velocity sensor into the ice allowing the thermal structure within the upper part of the ocean mixed layer to be characterized. AOFB data are transmitted daily by Iridium satellite and near realtime data are available on the flux buoy Web site (http://www.oc.nps.edu/stanton/fluxbuoy/). [9] IMB 2010A and ITP 38 were deployed about 10 m apart (this distance did not change over the 7 month drift), with AOFB 20 deployed about 60 m from the other buoys. Around the end of May, Web camera images show that a lead opened between the AOFB and the other instruments on the floe. There was some relative motion between the AOFB and other instruments until around the end of August, but all instruments remained within the field of view of the NOAA-PMEL Web camera. Time series of floe speed, internal ice temperature, and upper-ocean temperature and salinity from the IMB and ITP span the 2010 melt season, defined by the initiation and end of ice base ablation measured by the IMB as between 25 June and 7 September 2010 (Figure 2). Note that defining the beginning and end of melt is somewhat subjective. For example, major bottom melt started about 29 June and there was still some minor, much slower, bottom melt after 7 September. Figure 2Open in figure viewerPowerPoint (a) Ice mass balance buoy (IMB) snow and ice thickness and internal ice temperature. (b) ITP temperature difference from freezing at 7 m with IMB water-column temperature difference from freezing from the 19 thermistors between 200 and 380 cm depth below the top surface of the ice. Vertical red lines mark the start and end of the melt period, which we define by initiation and end of bottom ablation measured by the IMB. Depth-time sections of (c) ITP potential temperature and (d) salinity. (e) Drift speed V of the ice floe computed from ITP GPS positions. [10] ITP calibration procedures are described by Johnson et al. [2007] and R. Krishfield et al. (unpublished manuscript, 2008) (see http://www.whoi.edu/page.do?pid=23096). Predeployment laboratory-derived calibrations were adopted for all temperature and pressure data (postdeployment laboratory calibrations of two recovered ITPs documented temperature and pressure offsets after 2–3 years of 0.001°C to 0.002°C and around 1 dbar, respectively). These are taken as the uncertainties of the final ITP temperature and pressure data. Adjustments to the laboratory conductivity calibrations were derived and applied as detailed by Krishfield et al. (unpublished manuscript, 2008) to achieve consistency with recently acquired ship-based salinity estimates for the region. The resultant ITP salinity data have an uncertainty (relative to the ship data) of 0.005 or less. IMB thermistors were calibrated by applying an offset (specified by setting the initial thermistor value equal to freezing temperature based on the 7 m salinity of the ITP) to each thermistor series. The magnitude of the offset was no more than 0.06°C. This assumes a homogeneous mixed layer at freezing temperature on 20 April above the top depth sampled by the ITP. Freezing temperature was calculated using the ITP salinity time series at 7 m. IMB temperature measurements (between 200 and 380 cm below the top surface of the ice) do not deviate from ITP temperatures at 7 m depth (the top depth sampled by the ITP) over the duration of the drift. While this might be expected since ITP vertical profiles indicate mixed-layer depths were never shallower than about 12 m over the course of the IBO drift, it allows us to rule out that water properties immediately below the sea ice were different than those sampled by the ITP at 7 m. Summer Melt [11] Over the melt period, IMB 2010A documented a net sea ice bottom ablation of 0.40 m, and negligible surface ice melt. Total snowmelt was 0.35 m. The measured bottom ablation implies the total heat input to the ice cover during the melt period was about 119 MJ m−2 (given the latent heat of fusion 3.34 × 105 J kg−1, and assuming no increase in heat content of the ice base; Figure 2), equivalent to an average net ocean-to-ice heat flux over the melt period of about 18 W m−2. [12] The ocean-to-ice heat flux can be estimated by [McPhee, 1992] where cp = 3980 J kg−1 is the specific heat of seawater, cH = 0.0057 is a heat transfer coefficient [see McPhee et al., 2003], δT is the difference between mixed-layer temperature and freezing temperature (a function of mixed-layer salinity), and u*0 is the interface friction speed. The velocity u*0 may be estimated from ice-drift velocity (V, where, as is common, we assume the geostrophic flow is much smaller than the wind-driven ice-drift velocity) using a Rossby similarity relationship (see McPhee [2008] for a full discussion): where u*0 and V are horizontal vectors expressed as complex numbers, κ = 0.4 is von Karman's constant, and f is the Coriolis parameter with constants A = 2.12, B = 1.91. One source of error is associated with estimating the undersurface roughness z0 [see Wettlaufer, 1991]. Following McPhee et al. [2003], we take z0 = 0.01 m for consistency between our estimates and heat fluxes inferred from an IBO deployed in April 2002 as part of the NPEO [McPhee et al., 2003; see also Krishfield and Perovich, 2005]. Also following McPhee et al. [2003], we remove tidal and inertial components from V before applying (2) under the assumption that the inertial component of shear at the ice-ocean interface can be neglected and the ice and upper ocean react in the same way to tidal forcing (further discussion is given by McPhee [2008]). Employing (1) and (2) yields a mean ocean-to-ice heat flux over the melt period of about 15 W m−2 (Figure 3), roughly consistent with the observed sea ice bottom melt. Note that in this case the heat that causes basal ice melt derives entirely from surface-ocean warming by incoming solar radiation through open leads, rather than entrainment of heat from warmer water below the mixed layer; during the melt period the warm surface water is insulated by a cool halocline from the underlying warm water (Figure 2). The mean heat flux is not significantly different from heat flux estimates over the same period from the 2002 NPEO IBO analyzed by McPhee et al. [2003, Figure 3]. Figure 3Open in figure viewerPowerPoint (top) Interfacial friction speed from ice velocity and equation (2). (bottom) Estimated basal heat flux derived from equation (1). Vertical red lines mark the start and end of the melt period. [13] For comparison, interfacial-friction velocity was obtained directly by the correlation of the fluctuating horizontal and vertical components of velocity at 4 m measured by AOFB 20 (Figure 4). Variable u*0 measured from 40 min ensembles of Reynolds stresses was typically only about 2/3 the magnitude of u*0 derived using equation (2), suggesting that the value of z0 used to characterize undersurface roughness is too large for this floe, at least in the vicinity of the AOFB. Note that the two u*0 time series are only weakly correlated, with a correlation coefficient of 0.3, over the deployment. Over the melt period, the average heat flux calculated from equation (1) using the direct u*0 and the temperature series from the thermistor closest to the ice is about 10 W m−2; this heat flux would yield only about one-half of the observed bottom ablation (0.2 m compared to 0.4 m). Estimates of u*0 based on the Rossby similarity drag law depend on the choice of a roughness scale z0, while the direct measurements of u*0 at 4 m depth are sensitive to local under-ice morphology [see, e.g., Shaw et al., 2008]. The Web camera images show substantial heterogeneity of the ice surface over the course of the melt season that is likely replicated on the bottom surface of the ice. Further, the AOFB is positioned on the other side of a working lead for most of the melt period. We will show that even taking melt measured by the IMB as an upper bound for the region, the measured freshening (compared to previous years) is far too large to be attributed to melt. Figure 4Open in figure viewerPowerPoint (top) Arctic Ocean flux buoy (AOFB) temperature difference from freezing at eight levels (separated by 0.4 m) above approximately 4 m below the top surface of the ice. The top thermistor (warmest) indicates that the mixed layer was largely isothermal during the summer, with warmer surface waters around the beginning of July. (middle) Eddy-correlation-based interfacial-friction speed measured at 4 m below the ice (measured by AOFB 20). Note that the u*0 levels are lower than estimates in Figure 3, implying that the ice was smoother at the IBO site than inferred by the z0 = 0.01 m roughness value used in equation (2). (bottom) Estimated basal heat flux derived from the highest thermistor value (Figure 4, top), u*0 (Figure 4, middle), and equation (1). Vertical red lines mark the start and end of the melt period. [14] Given the densities of ice ρi ≈ 900 kg m−3 and snow ρs ≈ 330 kg m−3, the total melt is equivalent to a mixed-layer input of about 0.47 m of fresh water. This is an upper bound, assuming all the snowmelt equivalent fresh water (0.10 m) is input to the surface ocean (i.e., we do not account for snow sublimation [e.g., Déry and Yau, 2002] or collection of melt in melt ponds, both of which would reduce the equivalent freshwater input). For comparison, freshwater content of the surface mixed layer over the course of the IBO drift (relative to freshwater content on 25 June 2010 at the onset of the melt period) can be calculated assuming the mixed-layer salinity remains uniform from above the top depth (7 m) sampled by the ITP to the ice-ocean interface. If mixed-layer salinity on 25 June is S0, relative fresh water content (in meters) for a mixed layer of salinity S and depth D is (S0 − S)D/S0. Mixed-layer depth D is calculated to be where the potential density relative to 0 dbar first exceeds the shallowest sampled density by 0.01 kg m−3 [see Toole et al., 2010]. We estimate a change in freshwater content of the ocean mixed layer of about 0.7 m between 25 June and 7 September, about 30% more than can be attributed to melt (Figure 5). In general over the melt season, the relative freshwater content increases as the mixed layer thins; note the significantly larger impact of the change in salinity (compared to changes in mixed-layer depth) on freshwater content for the ranges of change observed in mixed-layer depth and salinity in this case. An alternative approach to calculating upper-ocean freshwater content is to integrate the salinity anomaly between a fixed depth and the ice-ocean interface. This approach, with a fixed depth of 25 m (the mean mixed-layer depth over the melt period), yields a similar freshwater content increase (about 0.7 m over the melt period). While it would appear that there is excess fresh water in the region that cannot be attributed entirely to seasonal melt, the possibility exists that spatial gradients sampled by a drifting ITP could be interpreted as temporal changes. However, as will be discussed in the next section, observations indicate that the entire region was significantly fresher in 2010 compared to 2007 and 2008 even though seasonal melt was comparable in all 3 years. Figure 5Open in figure viewerPowerPoint Change in freshwater content of the surface mixed layer over the course of the IBO drift (calculated relative to freshwater content on 25 June 2010 at the onset of the melt season) (solid curve). Mixed-layer depth defined as that point where the potential density relative to 0 dbar first exceeded the shallowest sampled density by 0.01 kg m−3 (shaded curve). The mixed-layer salinity was taken as the mean over the mixed layer and was assumed to remain uniform from above the top depth (7 m) sampled by the ITP. Vertical red lines mark the start and end of the melt period. Interannual Variability [15] Over the course of the 2010 melt season the IBO drifted about 500 km south, beginning in late June when it turned off the Lomononosov Ridge axis around 88°N. The fresher mixed-layer recorded during this drift is in contrast to surface salinity distributions of preceding years (Figure 6). ITP drifts in 2007 (ITP 7) and 2008 (ITP 19) indicated saltier mixed layers with correspondingly lower freshwater content in the vicinity of the 2010 IBO drift track, even though freshwater input due to melt in 2007 and 2008 was roughly comparable to that in 2010. An IMB was part of a manned station during the drift of the schooner Tara from the North Pole region toward Fram Strait over the 2007 melt period [see Nicolaus et al., 2010]. The IMB indicated an ice thickness decrease of 0.63 m (0.53 m surface melt and 0.10 m bottom) and total snowmelt of 0.18 m between June and August 2007. The following year, an IMB (2008E) deployed with ITP 19 recorded a total ice thickness decrease of about 0.60 m and total snowmelt of about 0.20 m between June and August 2008. This yields surface-ocean freshwater equivalent increases from melt of about 0.6 m over the 2007 and 2008 melt periods, comparable to freshwater input from melt inferred from the IMB drifting in the same region in 2010. There was not an ITP deployment in the North Pole to Fram Strait region in 2009, but measurements from a Naval Postgraduate School Arctic Ocean Flux Buoy (AOFB), which had a conductivity sensor at nominally 4 meters below the sea ice base, indicate fresher mixed layers (comparable to 2010) measured by the AOFB until around the beginning of June when it crossed 88°N (Figure 6c). This suggests that the region along the prime meridian was in transition in 2009; saltier mixed layers, similar to conditions in 2007–2008, were recorded over the southern portion of the 2009 AOFB drift. IMB 2009A deployed on the same floe with the AOFB transmitted data only until 5 August 2009. Over this time, it recorded a total snowmelt of 0.48 m and ice melt of about 0.10 m, equivalent to a freshwater input of about 0.23 m, although with about one month remaining in the melt season. Figure 6Open in figure viewerPowerPoint Map of salinity at 10 m from North Pole Environmental Observatory and Switchyard conductivity-temperature-depth stations and ITP profiles between 1 April and 1