3.1. Model Comparison With 129I Measurements in the 1990s
 To assess the accuracy of model simulations of 129I transport through the Arctic Ocean, a quasi-synoptic data set of129I measurements from the mid-1990s covering large parts of the central Arctic Ocean is compared to model results inFigures 2–5. The 129I data set is derived from the collection of water samples from icebreaker [Carmack et al., 1997; Smith et al., 1998] and nuclear submarine expeditions [Smith et al., 1999]. In 1995 and 1996 as part of the SCICEX program, the eastern Canadian, the Makarov and parts of the Eurasian Basins were sampled at three depth levels (corresponding to submarine standard operating depths); the Polar Mixed Layer, the halocline and the upper AWL at 59 m, 134 m and 240 m, respectively. The 129I results in Figures 2–4illustrate the lateral distribution of 129I about 25 years after initial releases into European waters and show the extent to which Atlantic-derived water had spread across the Arctic Ocean.
Figure 2. (a) Observed 129I concentrations (107 at/l) for surface mixed layer depths (0–59 m) for 1994–1996. (b) Simulated 129I distributions (107 at/l) at 59 m for September 1995 are in good agreement with 1994–1996 data sets above and illustrate the eastward spread of high 129I labeled, Atlantic-origin surface water across the Makarov Basin to the Chukchi Plateau.
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Figure 3. (a) Observed 129I concentrations (107 at/l) at a depth of 134 m for 1994–1996. (b) Simulated 129I (107 at/l) at134 m for September, 1995 conform to measurements above and show the eastward spread of Atlantic halocline water across the Mendeleyev Ridge into the Canada Basin.
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Figure 4. (a) Observed 129I concentrations (107 at/l) at a depth of 240 m for 1994–1996. (b) Simulated 129I (107 at/l) at a depth of 240 m (Sept., 1995) are in good agreement with measured values above and show the advance of high 129I labeled, upper AW into the Canada Basin.
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Figure 5. (a) Observed 129I concentrations (107 at/l) on a section from Alaska to Svalbard for 1994–1996. (b) Simulated 129I (107at/l) on Alaska-Svalbard section from 1995. Both panels similarly show cores of recirculating water of Atlantic origin labeled with high129I at mid depths.
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 The most remarkable feature revealed by those measurements was a strong front in the Polar Mixed Layer (0–59 m depth) crossing the Arctic from the western Chukchi Plateau to the northwestern Eurasian Basin, just north of Greenland (Figure 2). Both observations and model results, indicate that the Polar Mixed Layer of the southern and eastern Canadian Basin was largely uncontaminated by nuclear fuel reprocessing inputs of 129I. This water was derived from inflow through Bering Strait and was characterized by fallout 129I levels of the order of 5 × 107 at/l that prevailed in surface waters of the Pacific Ocean. Note that in the model simulation no fallout was applied as an additional source, leaving initial and Bering Strait inflow concentrations of 129I at zero. The location of the front between Atlantic and Pacific origin water and the measured concentrations of 129I that characterize this front are accurately simulated by the model. The front outlines the position of the TPD which extended from the Chukchi Sea across the North Pole to Fram Strait in 1995, approximately following the 0°–180° axis of longitude. Some of the more subtle features of the simulation cannot be resolved with the observational data, although this may simply reflect the fact that the timing of these measurements did not coincide with the development of rapidly changing flow structures.
 By 1995, 129I concentrations in the surface mixed layers on the continental shelves of the Barents, Kara and Laptev Seas had increased to values in excess of 100 × 107 at/l (Figure 2a) [Josefsson, 1998; Smith et al., 1998; Raisbeck and Yiou, 2002] following the enhanced inputs from the European reprocessing facilities in the early 1990s (Figure 1, inset). This observational feature is clearly delineated in the model results (Figure 2b). The relatively low 129I concentrations (<50 × 107 at/l) in the southern Nansen Basin in both the observational (Figure 2a) and model (Figure 2b) results show that surface water flow directly into the Eurasian Basin from the Barents Sea is minimal and that the dominant flow of surface water is eastward through the chain of Russian shelf seas.
 At a depth of 134 m in the interior halocline, 129I concentrations were lower than those in the mixed layer for both measured (Figure 3a) and model results (Figure 3b). The front separating the low 129I levels in Pacific-origin water and the elevated129I levels (>50 × 107at/l) in Atlantic-origin water had been displaced eastward toward the Canada Basin compared to its position at a depth of 59 m. The highest129I levels (>100 × 107 at/l) were observed over the southern Makarov Basin and Chukchi Plateau. The simulated 129I concentrations for 1995 were marginally lower than the observations in this area, which were made over the period of 1994–1996. Differences between the model simulation and the observations are to be expected, given the large horizontal and vertical gradients in the 129I distributions in this region and the relatively low spatial and temporal sampling densities. The model results show higher 129I values for upstream regions over the Siberian continental slope, north of the Laptev Sea where the Lomonosov Ridge joins the continental slope. This high 129I water spreads into the southern Makarov Basin and along the Lomonosov Ridge about a year later, thereby bringing model and experimental 129I results into good agreement. This suggests that one cause for the relatively small differences between the model and experimental results may be a slightly slower advection in the model as compared to that which actually prevailed during this time period. Observations and model results both show the presence of a flow branch carrying radionuclides into the interior of the Canada Basin from the northern edge of the Chukchi Plateau, as previously noted in several tracer studies [Smith et al., 1999; Smethie et al., 2000].
 At a water depth of 240 m, corresponding to upper AWL, the eastward progression of the 129I tracer distribution is very pronounced in both the observations and simulation (Figures 4a and 4b). The plume of high 129I concentrations extends in southern and northern branches around the Chukchi Plateau with similar tracer levels and frontal progression evident in model results and observations. High 129I concentrations along the Lomonosov Ridge outline the AW boundary current separating from the continental slope north of the Laptev Sea. Model results and observations at 240 m also show the reservoir of uncontaminated Atlantic Water water having low levels 129I (<1 × 107 at/l) in the eastern Canada Basin, north of the Canadian Archipelago. Observations indicate a similar behavior from two stations south of the Alpha Ridge. During the 1990s, the simulation shows that this water was undergoing cyclonic recirculation along the Canadian side of the Alpha Ridge and past Greenland toward Fram Strait as previously inferred from mean age calculations based on applications of transit time distributions (TTDs) to 129I and CFC transient tracer measurements [Smith et al., 1999; Smethie et al., 2000; Tanhua et al., 2009].
 An 129I section from Alaska to Svalbard through the Canada Basin based on results from submarine and icebreaker missions in 1994–1996 [Smith et al., 2011] is compared to the model 129I section for 1995 in Figure 5. High 129I concentrations in the upper 200 m in the Eurasian Basin reflect the return flow of Atlantic-origin surface and halocline water toward Fram Strait within the TPD. The model simulation also shows a core of water entering with Fram Strait Branch Water (FSBW) on the northern slope of Svalbard centered at 300–500 m depth, which carries elevated129I concentrations (>50 × 107 at/l) associated with the early 1990s increase in the 129I input function (Figure 1, inset). Lower concentrations over the Lomonosov Ridge at FSBW depths indicate that the early 1990s' 129I increase had not yet arrived at the North Pole by 1995. As noted in the previous paragraph both observed and model results show extremely low 129I levels (<1 × 107 at/l) in intermediate water in the interior of the Canada Basin south of the Alpha Ridge, a feature associated with the high mean age of this water mass [Smith et al., 1999; Smethie et al., 2000]. Both the core of the AW branch separating from the boundary current north of the Chukchi Plateau and ventilating the interior of the Canada Basin and the downstream core of the boundary current flowing eastward along the continental slope are evident as local maxima with 129I concentrations of 20–40 × 107 at/l in both the measured and simulated sections. The observations and model simulation also indicate the presence of a second 129I maximum in the lower AWL (500 m–1000 m) over the Eurasian Basin flank of the Lomonosov Ridge. The horizontal pattern of 129I concentrations at 600 m depth (Figure 6) indicates that the source for this 129I maximum is Barents Sea Branch Water (BSBW) which has entered the southern Nansen Basin via the St. Anna Trough and separated from the AW boundary current north of the Laptev Sea. This feature is replicated in the model simulation by an 129I maximum of similar magnitude (30–40 × 107 at/l) to observed values centered at about 1000 m. Generally, the concentrations and structural features of the simulated 129I distributions are close to those of the measurements. One important difference occurs at depths below 1000 m in the Canada Basin, where measured 129I levels of 5–10 × 107 at/l are higher than model values of 0–5 × 107 at/l. It is possible that the observed values at this depth carried a contribution from fallout or other unidentified sources from earlier time periods which have not been included in the model input function. It is also possible that the magnitude of dense water formation in the model is insufficient to simulate actual 129I concentrations below 1000 m. In principle this latter possibility can be examined using additional tracers.
Figure 6. Simulated 129I concentrations (107 at/l) at a depth of 600 m for September, 1995 show elevated levels extending along the continental slope to the Barents Sea indicating that 129I maximum at 600–800 m in Eurasian Basin (Figure 5) is associated with Barents Sea Branch Water (BSBW).
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 In summary, the three dimensional distribution of 129I in the Arctic Ocean basins, as characterized by quasi-synoptic measurements in the mid-1990s, is well described by the model simulation. Given that the129I input function had a strong increase in the early 1990s, these results also indicate that the advective time scales for the model circulation are realistic within an uncertainty of about one year.
3.3. Interruption of the AWL Boundary Current
 Two different flow features began to develop simultaneously in the mid-2000s: the AW boundary current flowing eastward from the Eurasian Basin via the Makarov Basin along the slope of the Chukchi Sea into the Canada Basin began to weaken with a corresponding strengthening of the return flow toward Fram Strait along the Mendeleyev Ridge (Figures 9b and 9c). This flow pattern is in contrast to the dominant pattern of the 1990s and early 2000s, when AW flow crossing the Lomonosov and Mendeleyev Ridges into the Canada Basin and the return flow along the central ridges were of similar strength (Figure 9a). This change in circulation regime was distinguished by a withdrawal of high 129I, labeled AW from the Alaskan slope to the Mendeleyev Ridge (compare Figures 9a and 9b). The period of intense AW boundary current transport and high eastward flow velocities on the Alaskan slope, which prevailed in the 1990s and early 2000s, had come to an end by the mid-2000s. The flow of upper AW (200–250 m) actually switched to a westward direction in 2004 (Figure 13, top, light green line) in concert with an intensification of the anti-cyclonic circulation of the Beaufort Gyre in the late 2000s which was characterized by a strong negative anomaly of the curl of the surface velocity (Figure 13, top, black line). At deeper levels the Atlantic Water boundary current at the Alaskan slope continued to circulate eastward, but at strongly reduced velocities compared to those typical of the 1990s (Figure 13, top, red and black lines). The change in circulation regime in the late 2000s, as is evident from the changes in the simulated tracer distribution in the Canada Basin and velocity changes at the Alaskan slope, was associated with the strengthening of the Beaufort Gyre and a deepening of the surface mixed layer, together with an increase in its freshwater content [McPhee et al., 2009; Proshutinsky et al., 2009; Rabe et al., 2011]. These changes resulted from a period of increased Ekman-pumping due to a negative anomaly of the wind + ice stress curl in the entire Canadian Basin starting in 2004 (Figure 13, top) [see also Rabe et al., 2011]. For the same region Proshutinsky et al.  found a doubling of negative wind stress curl for the 2000s as compared to earlier decades. Yang  suggested that an increase in Ekman pumping in the Beaufort Sea for the period 1998 to 2004 (with peaks in 1998 and 2003/2004) was a consequence of varying ice velocities associated with changes in ice dynamics (thinner and less areal coverage). His analysis covers the period 1979–2006 in the center of Beaufort Sea. Karcher et al. showed that anti-cyclonic flow of AW could occur in the Canadian and Makarov basin as a result of an artificially intense, anti-cyclonic Beaufort Gyre forced over a period of several decades.
Figure 13. (top) Ekman Pumping (cm/d, black curve) and curl of the surface currents (1/s, blue curve) in the Beaufort Gyre area (130° to 170°W, 70.5° to 80.5°N, for water depths > 500 m). The latter serves as a measure of the Gyre's intensity at the surface. (bottom) Mean eastward velocity (cm/s) at the Alaskan slope of the Beaufort Sea at water depths of 200–250 m (green curve), 310–385 m (red curve) and 480–580 m (black curve). All data are yearly means. The gray shaded years show periods of strong Ekman Pumping, which coincide with reduced eastward flow on the Alaskan slope.
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 The present work indicates that anti-cyclonic flow is possible for upper AW in a realistically forced setting. The shallower parts of the AW boundary current at the Alaskan slope seem to be more subject to anticyclonic flow reversal than the deeper layers. The results also show that previous periods of intensified Ekman pumping around 1960 and 1980 coincided with an intense Beaufort Gyre circulation and the reduction or even reversal in the AW boundary current at the Alaskan slope at all intermediate depth levels (Figure 13, gray shading). It should be noted that in the model simulation the direction of flow at the Alaskan slope is not necessarily representative of the circulation in the entire Canada Basin. Instead, it may be part of a smaller anticyclonic recirculation cell in the Beaufort Sea while the overall AW circulation in the Amerasian Basin may continue to be cyclonic, but with a shortened diversion extending from the Chukchi Plateau to the Canadian slope. Such situations occur in the model simulation in the 1950s to the 1970s (not shown). Early work from Newton and Coachman , based on Coachman and Barnes , found indications for such anticyclonic motion of AW in the Beaufort Sea, and a shortcut of AWL flow as described above. Their work is based on sparse data mostly from the 1960s, and deduces the flow direction based on the retention of AW characteristics, while disregarding temporal variability in the hydrography.
 The different hypothesized circulation regimes for the 1980s, 1990s and the late 2000s are illustrated schematically in Figure 14 for surface water (Figures 14a, 14b, and 14c) and AWL depths (250–350 m depth; Figures 14d, 14e, and 14f). An anti-cyclonic regime prevailed at the surface in the 1980s (Figure 14a), which nevertheless allowed cyclonic flow at AWL depths (Figure 14d). It developed into a weaker anti-cyclonic surface circulation with a smaller Beaufort Gyre and a TPD relocated toward North America in the 1990s (Figure 14b). At AWL depths, during the 1990s intense cyclonic flow occurred in the Canadian and Makarov Basins, supported by strong AW flow across the Lomonosov Ridge (indicated by doubled arrows in Figure 14e). After 2004, the intense anti-cyclonic circulation of an enlarged Beaufort Gyre dominated in surface waters (Figure 14c). There was also a variable westward flow of water with an elevated fraction of Atlantic origin water separating from the Transpolar Drift north of Greenland and an accompanying redirection of the flow of Pacific-origin water from the western part of Fram Strait into the Canadian Archipelago (Figure 14c). The cyclonic flow of the underlying AWL also underwent a strong reduction in intensity after 2004 with the weakened flow along the Lomonosov Ridge indicated by a dashed line in Figure 14f. At the upper AW level a switch to large scale anti-cyclonic flow in the Canada Basin may have occurred, resulting in the development of a third type of circulation regime by the late 2000s (Figure 14f). However, the resulting flow around the Chukchi Plateau is not entirely clear from the present state of knowledge; thus this entire regime in this area has been denoted by a question mark in Figure 14f. Undoubtedly the flow of water north of the Chukchi Plateau into the central Canadian Basin originally stems from the boundary current, but it is unclear whether it has separated from the return flow along the Alpha-Mendeleyev Ridge or it separated from the northern slope of the Chukchi Plateau (Figure 14f, black arrow). Following 2004, intermediate waters in the Canada and Eurasian Basins appear to have lost their direct connection via the boundary currents formerly flowing eastward through the Makarov Basin over the Siberian continental slope. Instead, flow in the Makarov Basin has become more directly linked with the cyclonic circulation regime of the Eurasian Basin. Such a switch of regimes of the Makarov Basin has previously only been known to occur at the surface.
Figure 14. Circulation pathways for Atlantic-origin (red lines) and Pacific-origin (black lines) water in the Arctic Ocean. (a) Surface flow in 1980s, (b) surface flow in 1990s, (c) surface flow after 2004 (dashed lines denote intermittent flow), (d) Atlantic Water Layer (AWL) flow in the 1980s, (e) AWL flow in the 1990s (double arrows in southern Makarov and Canada Basins indicate pathway for accelerated boundary current flow), and (f) AWL flow after 2004. Double arrows in southern Makarov Basin and along Mendeleyev Ridge indicate pathways for accelerated AW flow. The dashed line along Lomonosov Ridge indicates weakened AW flow. The flow in the Chukchi Plateau area (dotted) is less clear and thus tagged with a question mark. North of the Chukchi Plateau high129I water is found in the deep Canadian Basin (black arrow). This figure illustrates the shift from anticyclonic to cyclonic circulation regimes between the 1980s and 1990s and a return to an anticyclonic regime after 2004 accompanied by flow reversal for upper AW in the Canada Basin.
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 These results are consistent with observations of McLaughlin et al. of a strong, mid-depth transport of AW along the continental slope south of the Chukchi Plateau in the late 1990s combined with the diminished propagation of anomalously warm, Atlantic boundary current water on the Alaskan slope by 2005–7. Our results are also consistent with the observations of diminished anti-cyclonic residual transport of AW along the slope in the Beaufort Sea in the early 2000s [Nikolopoulos et al., 2009]. McLaughlin et al. also speculate about anti-cyclonic, interior circulation in the Canadian basin, as inferred from dynamic height results at 400 m (ref. to 1000 m) based on 2003–2007 data. The presence of weak, anti-cyclonic flow in the interior of the Canada Basin, including an eastward off-slope transport of AW from the northern tip of the Chukchi Plateau and a general southward flow tendency in the eastern Canada Basin provides additional support for the present model results. One further feature of the simulation that can be tested is an intermittent surface flow in the late 2000s (indicated by dashed lines inFigure 14c) which carries Atlantic-origin water westward along the northern slope of Greenland toward the Canada Basin leading to an intermittent blockage of the southward flow of Pacific-origin water toward Fram Strait. These model results are supported by observations of elevated129I levels (>200 × 107 at/l) in water collected north of Greenland in 2009 (Figure 8a) that clearly label this water mass as being mainly of Atlantic origin. Support for this westward flow is also provided by geostrophic velocities at depths of 50–60 m (relative to 500 dbar) for 2008 [Morison et al., 2012].