# Seasonal and Decadal Geostrophic Pathways of Pacific and Atlantic Waters in the Arctic Amerasian Basin From Observations

Noémie Planat, L. Bruno Tremblay, Carolina Dufour, David Straub

► To cite this version:

Noémie Planat, L. Bruno Tremblay, Carolina Dufour, David Straub. Seasonal and Decadal Geostrophic Pathways of Pacific and Atlantic Waters in the Arctic Amerasian Basin From Observations. Journal of Geophysical Research. Oceans, 2025, 130 (3), e2024JC021560 (22p.). 10.1029/2024JC021560. [hal-05028560](https:// HAL Id: hal-05028560 )

https://hal.science/hal-05028560v1

Submitted on 11 Apr 2025

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Distributed under a Creative Commons Attribution 4.0 International License
JGR Oceans

RESEARCH ARTICLE

10.1029/2024JC021560

**Special Collection:**

The Arctic Ocean's changing
Beaufort Gyre

**Key Points:**

* The time mean and seasonal Arctic circulations are visualized from two climatologies using Montgomery potentials on three isopycnals layers

* A strong anticyclonic Beaufort Gyre is detected in the layer of the summer Pacific Waters and vanishes into the layer of Atlantic Waters

* Waters from the Eurasian shelf outflow at Fram and Nares Straits, whereas that from the Pacific exit through Amundsen and M'Clure Straits

**Supporting Information:**

Supporting Information may be found in
the online version of this article.

**Correspondence to:**

N. Planat,
noemie.planat@mail.mcgill.ca

**Citation:**

Planat, N., Tremblay, L. B., Dufour, C. O., & Straub, D. (2025). Seasonal and decadal geostrophic pathways of Pacific and Atlantic Waters in the Arctic Amerasian Basin from observations. *Journal of Geophysical Research: Oceans, 130*, e2024JC021560. https://doi.org/10.1029/2024JC021560

Received 12 JUL 2024
Accepted 20 FEB 2025

**Author Contributions:**

**Conceptualization:** Noémie Planat,
L. Bruno Tremblay, Carolina O. Dufour,
David Straub

**Formal analysis:** Noémie Planat

**Investigation:** Noémie Planat

**Methodology:** Noémie Planat,
David Straub

**Project administration:** Noémie Planat

**Supervision:** L. Bruno Tremblay,
Carolina O. Dufour, David Straub

**Writing – original draft:** Noémie Planat

© 2025. The Author(s).
This is an open access article under the
terms of the Creative Commons
Attribution License, which permits use,
distribution and reproduction in any
medium, provided the original work is
properly cited.

<figure><img src="image_2.png" /></figure>

# Seasonal and Decadal Geostrophic Pathways of Pacific and Atlantic Waters in the Arctic Amerasian Basin From Observations

Noémie Planat¹ ID, L. Bruno Tremblay¹ ID, Carolina O. Dufour¹,² ID, and David Straub¹ ID

¹Department of Atmospheric and Oceanic Sciences, McGill University, Montreal, QC, Canada, ²University Brest, CNRS, IRD, Ifremer, Laboratoire d'Océanographie Physique et Spatiale (LOPS), IUEM, Brest, France

**Abstract** We examine the geostrophic pathways of Pacific and Atlantic Waters along isopycnal surfaces of the Arctic Ocean from Montgomery potentials using two observation-based climatologies: the World Ocean Atlas and the Monthly Isopycnal & Mixed-layer Ocean Climatology. The decadal mean circulation (2005–2017) shows an anticyclonic circulation for both the summer Pacific Waters (sPW) and winter Pacific Waters (wPW, although somewhat weaker) in the Canada Basin with subduction along the Chukchi plateau for both water masses. On the wPW layer, the diagnostics also highlight a year-long persistent flow from Herald Canyon on Chukchi Shelf to the Canadian Arctic Archipelago around the Canada Basin. Outflows are found at Nares and Fram Straits at the density of sPW and wPW. These outflowing waters, previously reported as Pacific Waters, are found here to originate from the Eurasian shelf. Deeper, the Atlantic Waters flow anticyclonically around Northwind Ridge and cyclonically along the Alaskan shelfbreak forming a boundary current. The subduction of warm water along the Alaskan and Chukchi shelves is found to occur in summer on the sPW isopycnal layer, thus contributing to transferring heat into the center of the gyre. However, an advection of virtual Lagrangian particles from Pt. Barrow suggests that processes operating on timescales shorter than a month or at interannual timescales are the main contributors to the heat build up observed over the past two decades. This analysis supports the hypothesis that high frequency processes are key in shaping the subsurface heat reservoir of the Canada Basin.

**Plain Language Summary**
Two heat reservoirs lie below the sea ice cover in the Canada Basin of the Arctic Ocean—the summer Pacific and Atlantic Waters. Here, we focus on the trajectories of these warm waters upon entrance in the Canada Basin for the period 2005–2017 using two different observational data sets and based on the observation that water flows along layers of constant density. Results show a well-defined anticyclonic circulation of the sPW at 60–100 m depth; a weak and anticyclonic circulation of the winter Pacific Waters at 100–200 m depth and a cyclonic circulation of the Atlantic Waters layer deeper between 200 and 700 m. The analysis of the upper layers pathways also highlight an important outflow of waters originating from the Eurasian shelf through Nares and Fram Strait, previously identified as Pacific Waters. Seasonal analysis of these pathways highlights large variability in the location and shape of the gyre, with important subduction of Pacific Waters along Chukchi Sea. An advection of particles along these geostrophic pathways demonstrate the need to account for small-scale processes (eddies, filaments, etc.) and interannual variability to explain the build-up of the heat reservoir that has occurred at the centre of the Beaufort Gyre over the past decades.

# 1. Introduction

## 1.1. Context and Motivation

The Canada Basin of the western Arctic is known to have a singular vertical hydrography in which two heat reservoirs reside distinctly on top of each other and below the sea ice. These correspond to the summer Pacific Waters (sPW ≈ 60–100 m) and the Atlantic Waters (AW ≈ 250–600 m), and are separated vertically by a layer of cold winter Pacific Waters (wPW ≈ 100–200 m). In autumn and early winter, additional remnant heat from previous summer is also trapped beneath the seasonal pycnocline at around 30 m, with this heat being brought up into the mixed layer by wind mixing and brine-induced convection later in the season. The two more permanent heat reservoirs have the potential to greatly impact the sea ice cover should the heat they contain be brought to the surface (Ladd et al., 2016; Polyakov et al., 2010; Rudels et al., 2004; Steele & Boyd, 1998; Timmermans & Toole, 2023; Turner, 2010). Helping to prevent this are two successive haloclines, the upper and lower haloclines,
Writing – review & editing:

L. Bruno Tremblay, Carolina O. Dufour,
David Straub

which insulate the surface mixed layer and sea ice from this heat (Aagaard et al., 1981; Rudels et al., 1996). Indeed, the upward heat flux from the AW layer, achieved through double stair-case diffusive processes, is found to be small in the Canada Basin (circa 0 – 0.3 W/m²; Lique et al., 2014; Timmermans et al., 2008). It is interesting then that the recent melt of multi-year sea ice in the Western Arctic—observed despite a positive trend in multi year ice advection into the Beaufort Sea (Babb et al., 2022)—is collocated with these two subsurface heat reservoirs. This has led to the suggestion that, despite insulation provided by the upper halocline, heat from the sPW may have contributed significantly to this melt (Woodgate et al., 2010). Related to this suspicion is that the upper sPW has been observed to warm and shallow over the past decade (Rosenblum et al., 2021). Within the marginal seas, where most rapid sea ice decline was recently documented along with an increase in solar heat input (Perovich et al., 2007), the sPW are found to be important for initiating the ice-albedo feedback (Woodgate et al., 2010). Understanding the penetration of the PW and AW in the Beaufort Sea is thus key to determine the build up of the subsurface reservoirs, and hence to project the future evolution of the sea ice in the Arctic.

**1.2. Pacific Waters**

The Pacific Waters (PW) enter the Arctic through Bering Strait, with a mean inflow estimate around 1 Sv (Woodgate et al., 2005). They reside on the Chukchi shelf for about five to 9 months. There, they acquire distinct properties depending on the season (Panteleev et al., 2010), thus resulting in the wPW and the sPW. As wPW flow on the shallow, ice-covered shelf during winter months, they become colder, saltier, and more nutrient-rich (Codispoti et al., 2005). Conversely, sPW are formed on the shelf during summer months, when ice disappears, temperatures warm, and salt and nutrient levels decrease (Tremblay et al., 2015). The pathways of both wPW and sPW from the Chukchi shelf to the Canada Basin are relatively well understood, with the three main entrance points being Barrow Canyon (East), the Central Channel (middle) and Herald Valley (West, e.g., Gong & Pickart, 2015; Ladd et al., 2016; Pickart, 2004; Spall et al., 2018; Weingartner et al., 2005). A secondary pathway through Long Strait Shelf to the East Siberian Shelf has also been suggested from model studies (Lique et al., 2010; Nguyen et al., 2011).

Beyond the shelf, the fate of these waters is less certain. Some evidence suggests that they are in part deflected eastward, feeding the Alaskan Coastal Waters along the shelfbreak and potentially reaching the Canadian Arctic Archipelago and Nares Strait (Brugler et al., 2014; Hu & Myers, 2013; Nikolopoulos et al., 2009; Pickart, 2004; Steele et al., 2004; Timmermans et al., 2017). Another part is thought to follow the Chukchi shelf westward, eventually turning North and entering the Central Arctic along the Mendeleev Ridge (Jones et al., 1998; Steele et al., 2004) before circulating anticyclonically within the Canada Basin, forming the so-called Beaufort Gyre. Both branches are baroclinically unstable and are thus prone to the generation of mesoscale eddies (Hunkins, 1974; Spall et al., 2008; von Appen et al., 2022), which in turn participate in the transport of heat and salt toward the Canada Basin (MacKinnon et al., 2021; Spall et al., 2008). Such eddies are ubiquitous in observations and high-resolution models along the shelfbreak, as well as in the basin (e.g., Kubryakov et al., 2021; Manucharyan et al., 2022; Wang et al., 2020; Zhao et al., 2014), although uncertainties remain regarding the precise mechanisms controlling their formation and evolution (Manucharyan & Stewart, 2022). Flow of PW along these two branches is thought to subduct at locations following the seasonal migration of isopycnals (Timmermans et al., 2014, 2017). They further suggest a helical pathway into the gyre interior for this subducted water. Overall, the relative importance of eddy transport and subduction in feeding the Canada Basin with PW remains uncertain, as does the precise pathways and partition of these waters once in the Arctic interior.

In the Canada Basin, the circulation (and therefore, PW pathways) is forced mainly by winds related to the semi-
permanent atmospheric Beaufort High, although this forcing is modified considerably by the ice cover (Thorndike
& Colony, 1982). The transfer of momentum from the atmosphere to the upper ocean is most efficient in fall,
when the winds are strong and the ice weak (Brenner et al., 2021; Cole et al., 2017; Martin et al., 2014), whereas in
late winter, thicker and more compact ice pack tend to reduce the momentum transfer and act as a drag on the
upper anticyclonic oceanic geostrophic flow (the so-called ice-ocean governor; e.g., Meneghello et al., 2017). The
response of the upper ocean to changes in wind forcing and sea ice circulation is thought to occur on the order of a
few months to a few years (Doddridge et al., 2019). Variability in the PW circulation is therefore affected by
seasonal to decadal variability in the Beaufort High (e.g., the Arctic Oscillation or Arctic Ocean Circulation;
Proshutinsky & Johnson, 1997; Thompson & Wallace, 1998) as well as changes in sea ice thickness and con-
centration (Kwok, 2018).
### 1.3. Atlantic Waters

The Atlantic Waters (AW) enter the Arctic through two main gateways: the Fram Strait and the Barents Sea Opening (e.g., Rudels et al., 2004). These two gateways account for most of the heat input to the Arctic, with contributions of 50–70 TW from the Barents Sea Opening and 30–42 TW from the Fram Strait (Beszczynska-Möller et al., 2011) for a total heat transport to the Arctic estimated around 91–167 TW (Beszczynska-Möller et al., 2011; Cuny et al., 2005).

The AW entering through Fram Strait follow the continental shelf cyclonically at 150–600 m depth in the Nansen Basin (Pérez-Hernández et al., 2019) and are joined by the Barents Sea branch at St. Anna Trough. This branch is cooler and fresher, and lies below the Fram Straight branch, at around 700 m (McLaughlin et al., 1996; Schauer et al., 2002). Both branches continue cyclonically in the Nansen basin (Schauer et al., 2002), roughly following the shelfbreak, and the western sides of Gakkel and Lomonosov Ridges in the central Arctic (Dmitrenko et al., 2008; Holloway et al., 2007; Isachsen et al., 2003; Ivanov et al., 2009; Timmermans & Marshall, 2020; Woodgate et al., 2001). Beyond Lomonosov Ridge, part of the flow is believed to continue along the Eurasian shelfbreak (~3 Sv, Woodgate et al., 2001) but little is known on the exact path around the Northwind Ridge (Karcher et al., 2012; Woodgate et al., 2007). Possible pathways reported are along the Alpha-Mendeleev Ridge (Rudels et al., 2013), around the Northwind Ridge (both cyclonically and anticyclonically, Shimada et al., 2005; McLaughlin et al., 2009; Woodgate et al., 2007) and above the Northwind Ridge directly into the Canada Basin (Smith et al., 1999). In the Canada Basin, observations report departures from the cyclonic boundary current in the form of double boundary currents (Li et al., 2020), intrusions in the Canada Basin (Carmack et al., 2013; May & Kelley, 2001; Woodgate et al., 2007) or entrainments from the above anticyclonic flow (McLaughlin et al., 2009).

This cyclonic circulation of AW is controlled both locally by the upper ocean and ice circulation, and remotely by the inflow from the Nordic Seas, the potential vorticity inflow at St. Anna Trough and the dense water formation over the Barents Sea shelf (Karcher et al., 2007; Schauer et al., 2002). These remote forcings are thought to follow the Nordic Seas decadal variability (Karcher et al., 2007; Polyakov et al., 2004, 2005) as well as decadal climate trends (Richards et al., 2022), whereas seasonal variability is negligible in the Amerasian Basin (Lique & Steele, 2012). Direct dynamical influence of the upper layer onto the AW layer might constrain the lateral extent of the boundary current along the shelfbreak (Lique & Johnson, 2015; Lique et al., 2015) although the coupling efficiency and direction between the two layers in the Canada Basin (from upper layers to AW vs. from AW to the upper layers) remains unquantified.

### 1.4. Objectives and Outline

The pathways of the PW and AW contribute to build the complex stratification of the Canada Basin. Despite the increased number of observations in recent decades due to major observational programs such as the Ice-Tethered Profilers (ITPs) since 2004, MOSAIC expedition in 2019–2020, the Beaufort Gyre Observing System (BGOS) and Joint Ocean Ice Study (JOIS) expeditions since 2003, in situ measurements remain sparse. In addition to historically sparse observations, the computational prohibitive cost of running numerical simulations with sufficient horizontal and vertical resolution to resolve the complex vertical stratification and mesoscale processes (3 – 4 km in the Canada Basin; Hallberg, 2013; Nurser & Bacon, 2013) have also impeded our investigation of these pathways. Because of this, the mean pathways and variability of PW and AW as they leave the shelf and enter the Canadian and Amerasian Basins remain poorly known. However, recent observational programs have contributed to reach a higher density of observations in particular in the upper 700 m with ITPs that make observational climatologies more reliable in the most recent decade.

In this paper, and despite the uncertainties related to insufficient data coverage, we attempt to quantify geostrophic flow of PW and AW along isopycnal layers on climatological and seasonal timescales. Our main focus is on the shallower sPW, since these waters have the greatest potential to affect the sea ice balance in the near term. To this end, we use two climatologies that we describe in Section 2 and a dynamical diagnostic of the geostrophic flow on isopycnal layers, namely the Montgomery potential (e.g., McDougall, 1989, derived in Section 3). This approach follows the work done in the AW layer by Aksenov et al. (2011) and in the wPW by Zhong et al. (2019), a notable difference being that those studies were based on model output, whereas we use observational-based climatologies. Section 4.1 presents a description of the climatological geostrophic circulation of PW and AW as derived from the Montgomery Potential diagnostic provided for each layer and a discussion of the uncertainties and robustness of the method. Section 4.2 discusses seasonal variability of these pathways and
proposes a first order quantification of the “transient” flow needed for heat to enter the Beaufort Gyre on the sPW layer. Section 5 reports some key discussion points, as well as the main conclusions and proposed future work.

## 2. Data

### 2.1. Observation-Based Climatologies

We use the World Ocean Atlas (WOA), a climatology of monthly mean fields of practical salinity ($S_p$), in situ density ($\rho$) and in situ temperature ($T$, Locarnini, Boyer, et al., 2019; Locarnini, Mishonov, et al., 2019; Zweng et al., 2019). WOA builds on the World Ocean Database 2018 (Boyer et al., 2018). It includes about 9,000 profiles of density in the Canada Basin (red box on Figure 1) over the period 2005–2017. Detailed data coverage is displayed in Figures S4 (decadal) and S5 (seasonal at 70 m) in Supporting Information S1. Among these are all available vertical profiles from Ice-Tethered Profilers (ITP), processed to Level 3 of post-processing and covering the period 2004–2017 (Krishfield et al., 2014). All profiles between 2005 and 2017 are used to modify a background state which was derived from all available observations, including early records in the Canada Basin back from mid-20th century. The data coverage of decades prior to 2005 however remains too sparse to construct a robust climatology (Zweng et al., 2017). Errors on the temperature, conductivity, and pressure from the individual profiles of the World Ocean Database are of 0.001 – 0.005°C, 0.0003 – 0.002 S m⁻¹ and 0.015 – 0.08%, respectively (Boyer et al., 2018).

WOA has a 1/4° horizontal resolution and 57 levels in the vertical (e.g., with 5 m resolution over the first 100 m and about 50 m resolution at 1,000 m depth). The vertical coordinate, $z^\dagger$, is referenced to the surface and increases downward: $z^\dagger = \eta - z$, where $z = 0$ corresponds to the geoid and $\eta$ is sea level height (which is not corrected for the inverse barometer effect). We use WOA to assess the main circulation patterns in the Canada Basin during the period 2005–2017 and compare these results with the same analysis conducted on a second data set, the Monthly Isopycnal & Mixed-layer Ocean Climatology (MIMOC, Schmidtko et al., 2013). MIMOC has a similar resolution that is 1/2° horizontal resolution and 81 sea pressure vertical levels from the surface to 2,000 m, with a resolution of 5 dbar at 100 m. MIMOC is based on an earlier version of the World Ocean Database (version 2009, Boyer et al., 2009) and includes all ITPs available until 2011. It however presents strong biases in temperature along the Eurasian Shelf. MIMOC is built such that it is representative of the ocean state between 2007 and 2011, and thus is widely used to document the Arctic in the recent decade (e.g., Regan et al., 2020; Zhong et al., 2019). We assess the robustness of the key findings derived from WOA with this second data set in Section 4.1.

### 2.2. Sea Surface Height (SSH)

From the WOA and MIMOC data sets, we construct the thermal wind shear. To get the geostrophic circulation, an integration constant is needed, and this comes from Sea Surface Height (SSH). We use two different Sea Surface Height (SSH) data sets. The first (A-SSH) is the altimeter-derived product from Armitage et al. (2016). The initial (2004–2014, Envisat) and updated (2011–2020, CryoSat-2) versions have 2° longitude and 1/2° latitude spatial resolution and are corrected for tides and inverse barometer effects. As such they are directly comparable to SSH from ocean simulations that do not include tides. Because A-SSH does not provide SSH north of 81.5° N (Pole “hole”), we use a model-derived SSH data set (O-SSH) from the European Centre for Medium-Range Weather Forecasts (ECMWF) Oceanic Reanalysis System 5 (ORAS5 Zuo et al., 2019) to fill this gap. This reanalysis covers the full period of interest (2005–2017). O-SSH uses the ocean modeling platform NEMO version 3.4 with linear free surface and a 1/4° horizontal ORCA grid. The sea ice component is LIM2 run with levitating sea ice. It is forced by the atmospheric fields of the ECMWF Reanalysis version 5 (ERA5, Hersbach et al., 2020) and assimilates SSH between 50° S and 50° N from AVISO (Archiving Validation and Interpretation of Satellite Oceanographic data) DUACS (Data Unification and Altimeter Combination System, version DT2014 Pujol et al., 2016) for the period 1993 onward. It also assimilates passive microwave sea ice concentrations from the Operational Sea Surface Temperature and Ice Analysis (OSTIA) operational and re-processed data set, for the period 1979 onward (Donlon et al., 2012).

A linear sum of O-SSH and A-SSH is used to construct an hybrid SSH field in Section 3.3, namely W-SSH, which covers the entire Arctic, resembling A-SSH away from the pole and O-SSH over the pole. A comparison between A-SSH and O-SSH is available in Figure S6 in Supporting Information S1.
<figure><img src="image_3.png" /><figcaption>**Figure 1.** Bathymetry of the Arctic Ocean from the IBCAO V4.2.13 Grid product (Jakobsson et al., 2020). The green line indicates the location of a trans-Arctic transect referred to in Figures 2a and 2b. Following Lavoie et al. (2022), we define the Canada Basin as the area between 72 : 82°N, -163 : -130°W (red box). PPI refers to Prince Patrick Island. Gray lines indicate the 100 and 1,000 m isobaths.</figcaption></figure>

**3. Method**

**3.1. Identification of Water Masses From Potential Density**

In the following, we identify the sPW, wPW, and AW using density classes associated with key temperature features in the Canada Basin water column. This differs from previous studies (e.g., Athanase et al., 2020; Brugler et al., 2014; Coachman et al., 1975; Rudels et al., 1997; Steele et al., 2004; Timmermans & Marshall, 2020; Zhong et al., 2019), which used fixed salinity or density classes, or temperature extrema within fixed salinity/density bounds. The reason for the different method used here is that it can be repeated on numerical model output,
<figure><img src="image_4.png" /><figcaption>**Figure 2.** Conservative temperature (a) and absolute salinity (b) from WOA with surface referenced potential density (dashed lines) along the transect shown in Figure 1. (c) Mean vertical profile of conservative temperature and potential density referenced to surface in the Canada Basin in September from ITPs (all ITPs gathered between 2004 and 2017 at level 3 of post-processing). Different colors (schematically) refer to the five distinct water masses present in the Canada Basin. The solid circles and stars respectively indicate the temperature extremum and the corresponding potential density used to define the core of the water masses. The red box indicates the Canada Basin area.</figcaption></figure>

provided the models correctly simulate the vertical temperature structure, but despite the large biases they often have in salinity and therefore in density (e.g., Rosenblum et al., 2021).

For each data set, we identify the core of the sPW, wPW, and AW in the Canada Basin from, respectively, the local minimum (wPW) and local maxima (sPW and AW) of the spatially averaged decadal temperature profiles. The Canada Basin is defined as the region extending from 72 to 82°N, and from -163 to -130°W following Lavoie et al. (2022). We then extract the spatially and temporally averaged potential density value at the depth of these temperature extrema and assume the associated isopycnal to be representative of the three water masses across the Canada Basin (Figure 2c, Table 1). The three surfaces therefore represent the core of each water mass and not their full vertical extent, something to keep in mind for the AW layer, in particular, since it is relatively thick (see Figure 2c). The chosen isopycnals are consistent with previous observation-based studies that use

<caption>**Table 1**</caption>

Characteristic Potential Density ($\sigma_0$) and Absolute Salinity ($S_A$) of the sPW, wPW, and AW Cores in the Canada Basin in WOA, MIMOC, and the Ice-Tethered-Profilers (ITP) Data Set Determined From the Method Described in Section 3.1

<table><thead><tr><th></th><th>ITPs (2004–2017)<br>$\sigma_0$ [kg m<sup>-3</sup>]; min/max</th><th>WOA (2005–2017)<br>$\sigma_0$ [kg m<sup>-3</sup>]; min/max</th><th>MIMOC (2007–2011)<br>$\sigma_0$ [kg m<sup>-3</sup>]; min/max</th></tr></thead><tbody><tr><td>sPW</td><td>1,024.77;</td><td>1,024.90; 1,024.70 (Oct.)/1,025.05 (Mar.)</td><td>1,024.99; 1,024.57 (Sept.)/1,025.12 (June)</td></tr><tr><td>wPW</td><td>1,026.61;</td><td>1,026.50; 1,026.43 (Nov.)/1,026.76 (Jan.)</td><td>1,026.58; 1,026.52 (Jan.)/1,026.64 (Sept.)</td></tr><tr><td>AW</td><td>1,027.93;</td><td>1,027.92; 1,027.92 (Feb.)/1,027.93 (Jun.)</td><td>1,027.92; 1,027.92 (Nov.)/1,027.92 (Sept.)</td></tr></tbody><thead><tr><th></th><th>ITPs (2004–2017)<br>$S_A$ [g/kg]; min/max</th><th>WOA (2005–2017)<br>$S_A$ [g/kg]; min/max</th><th>MIMOC (2007–2011)<br>$S_A$ [g/kg]; min/max</th></tr></thead><tbody><tr><td>sPW</td><td>30.98</td><td>31.13; 30.89/31.32</td><td>31.24; 30.72/31.40</td></tr><tr><td>wPW</td><td>33.24</td><td>33.10; 33.01/33.41</td><td>33.19; 33.12/33.26</td></tr><tr><td>AW</td><td>34.99</td><td>34.98; 34.98/34.99</td><td>34.99; 34.98/34.99</td></tr></tbody></table>

Note. Minimum and maximum values are obtained from the full seasonal cycle.
different identification criteria (e.g., Brugler et al., 2014; Timmermans et al., 2014; Zhong et al., 2019). However, although some analysis report changes across time of the characteristic salinity or potential density of the water masses, we focus in this study on the decadal circulation and climatology. In that sense, the value of salinity that is used to represent each water mass in our analysis is only representative of the decadal averaged properties. Note that outside of the Canada Basin, water mass properties on these three isopycnals need not to correspond to the three water masses. However, for the sake of simplicity, we will refer to them as the “sPW, wPW, and AW layers.” Our results are robust to the exact choice of potential density defining the three layers. For instance, repeating the calculation with the potential density values at two grid points above or below the temperature extrema gives similar results for the circulation. Varying the depth of the temperature extrema by ±2 grid points on the vertical gives a range of ±Δσ of 0.4, 0.7, and 0.016 kg.m⁻³ for the sPW, wPW, and AW layers, respectively, which encompasses the small seasonality that exists for the depth of the sPW in our data sets. Our results are also robust to the choice of reference pressure for potential density of the AW layer (σ₁ vs. σ₀).

## 3.2. Circulation From Montgomery Potentials

To visualize the geostrophic circulation in the three layers, we use the Montgomery potential, M, to plot geostrophic streamlines on potential density surfaces. To this end, it is useful to introduce a Boussinesq version (P) of hydrostatic pressure (P<sub>H</sub>). P ignores compressibility effects, implying ∂<sub>z</sub>P = −σg, where σ is potential density and g is gravity. This differs from the actual hydrostatic pressure, which increases with depth in proportion to in situ density, ρ: ∂<sub>z</sub>P<sub>H</sub> = −ρg. Taking σ to be referenced to the surface, P<sub>H</sub> > P throughout the water column and at 1,000 m depth, this difference is about 10 hPa. Most of this, however, can be thought of as stemming from a reference compressibility profile that depends only on z. As such, horizontal derivatives are not affected. In other words, geostrophic velocities **u**<sub>g</sub> calculated using the two versions of pressure are approximately equal,

$$ \mathbf{u}_g = \frac{\hat{\mathbf{z}}}{\rho f} \times \nabla P_H \approx \frac{\hat{\mathbf{z}}}{\sigma f} \times \nabla P, \quad (1) $$

where f is the Coriolis parameter. A direct comparison of the two showed differences to be smaller than 10⁻⁴ m/s at 500 m on average (not shown).

The Montgomery potential can then be found by a change of vertical coordinate (from z to σ). In Cartesian coordinates, with **u**<sub>g</sub> = u<sub>g</sub>**x** + v<sub>g</sub>**y**,

$$ -fv_g = -\frac{1}{\sigma} \left. \frac{\partial P}{\partial x} \right|_z = -\left( \frac{1}{\sigma} \left. \frac{\partial P}{\partial x} \right|_\sigma + g \left. \frac{\partial z}{\partial x} \right|_\sigma \right) = -\left. \frac{\partial M}{\partial x} \right|_\sigma \quad (2) $$

$$ fu_g = -\frac{1}{\sigma} \left. \frac{\partial P}{\partial y} \right|_z = -\left( \frac{1}{\sigma} \left. \frac{\partial P}{\partial y} \right|_\sigma + g \left. \frac{\partial z}{\partial y} \right|_\sigma \right) = -\left. \frac{\partial M}{\partial y} \right|_\sigma $$

where σ and z(x, y, σ) correspond to the potential density and depth of the surface under consideration and the Montgomery potential is given as follows:

$$ M = \frac{P}{\sigma} + gz. \quad (3) $$

Note that, far below z = 0, M is a small difference of two large terms. With this in mind, it is convenient to rewrite it in a different form. First note that

$$ \begin{aligned} P &= P_{\text{atm}} + \int_{z}^{\eta} (\sigma_{c} + \delta\sigma(z')) g dz' \\ &= P_{\text{atm}} + \sigma_{c} g (\eta - z) + \int_{z}^{\eta} g \delta\sigma(z') dz', \end{aligned} \quad (4) $$

where the subscript on σ<sub>c</sub> is meant to emphasize that it is constant (i.e., σ<sub>c</sub> corresponds to the potential density surface used to derive the Montgomery potential M<sub>c</sub> and δσ = σ − σ<sub>c</sub>). Equation 3 thus reads:
<figure><img src="image_5.png" /><figcaption>**Figure 3.** Sea Surface Height from (a) Armitage et al. (2016, A-SSH), (b) the Ocean Reanalysis System 5 (O-SSH) and (c) a weighted sum of A-SSH and O-SSH (W-SSH; see Section 3.3 for details). The dashed black line on panel (a) marks the pole hole and the dashed blue/green lines on panel (c) are the latitudinal interval over which the weighted sum is calculated.</figcaption></figure>

$$M_c = \frac{P_{\text{atm}}}{\sigma_c} + g\eta + \frac{1}{\sigma_c} \int_z^{\eta} g\delta\sigma(z') \, dz'. \qquad (5)$$

Finally, defining SSH to be sea surface elevation ($\eta$), corrected for the inverse barometer effect,

$$\text{SSH} = \eta + \frac{P_{\text{atm}}}{\sigma_c g}, \qquad (6)$$

gives

$$M_c = g\text{SSH} + \frac{1}{\sigma_c} \int_z^{\eta} g\delta\sigma(z') \, dz'. \qquad (7)$$

In practice, the pressure and in situ density are computed iteratively from the practical salinity and in situ temperature, using the TEOS 10 equation for density (IOC et al., 2010), and the depth as an initialization of the pressure. Two iterations are sufficient for convergence to within $3 \cdot 10^{-5}\%$. For MIMOC, conservative temperature and absolute salinity are provided in pressure coordinates, so this step is unnecessary. Numerical integration following Equation 7 is then used to compute $M$ in z-coordinates (respectively P-coordinates for MIMOC) and later interpolated on the $\sigma_c$ surface.

### 3.3. Sea Surface Height Weighted Product (W-SSH)

Both O-SSH and A-SSH present similar patterns of SSH anomalies for the period 2011–2020 and have clear signatures of the Beaufort Gyre (Figures 3a and 3b). Important differences lie along the Chukchi shelfbreak, where the root mean square deviation reaches 15 cm (about 50% of the satellite-derived value, not shown). In the Canada Basin, both data sets co-vary similarly in time and amplitude (not shown). In particular, the important increase in SSH anomaly around 2008 reported in for example, Giles et al. (2012) is well represented by O-SSH with little spread between the five ensemble members. O-SSH is also in agreement with other reanalyzes (Lelouche et al., 2018) or observational SSH data sets (Doglioni et al., 2022).

At the scale of the Arctic Basin, the good agreement between O-SSH and A-SSH spatial patterns and temporal variability and the small spread within O-SSH ensemble members justifies the use of O-SSH ensemble mean to cover the Pole Hole. Over the Beaufort gyre region, however, O-SSH shows a stronger re-circulation and, along the Chukchi shelfbreak, differences between the two data sets can be as much as 50%. Because it is data-based, we take A-SSH as being closer to the truth and therefore construct a weighted product (W-SSH), which corresponds to A-SSH where satellite observations are available and transitions to O-SSH near and over the Pole Hole. Specifically, W-SSH is constructed as follows:

$$W_{\text{SSH}} = [1 - x(\lambda)] A_{\text{SSH}} + x(\lambda) O_{\text{SSH}} \qquad (8)$$
where $\lambda$ is the latitude and $x(\lambda)$ the weight function, defined as follows:

$$x(\lambda, \lambda_0, \epsilon) = \begin{cases} 0 & \text{if } \lambda < \lambda_0 - \epsilon \\ \frac{1}{2} \left[ 1 + \cos\left(\pi\left(\frac{\lambda - \lambda_0}{2\epsilon} - \frac{1}{2}\right)\right) \right] & \text{if } \lambda \in [\lambda_0 - \epsilon, \lambda_0 + \epsilon] \\ 1 & \text{if } \lambda > \lambda_0 + \epsilon \end{cases} \quad (9)$$

where $\lambda_0 = 78.5^{\circ}\text{N}$ is a tuning central latitude and $2\epsilon = 6^{\circ}$ is the width of the zone where W-SSH is a weighted sum, both chosen from the data sets' resolution. The resulting W-SSH resembles both products with data in the pole hole and a circulation that resembles that of A-SSH in the southern part of the domain (Figures 3a and 3c). Although the exact form of $x(\lambda)$ is arbitrary, it ensures a smooth transition between the two data sets ($x$ belongs to the class $C^1$), and is simple enough as it depends on only two parameters to control the width and latitude of the merge. The circulation presented in the rest of the paper is robust to the choice of $x(\lambda)$ (not shown).

# 4. Results

We begin with a brief description of atmospheric driving patterns, sea ice, and SSH fields, before discussing time mean pathways of the three water masses in the Canada Basin. Atmospheric forcing in the Canada Basin is dominated by a semi-permanent winter anticyclonic circulation, known as the Beaufort High (Figure 4a). It extends over the Amerasian Basin and is associated with a North–South asymmetry with strong easterly winds along the Alaskan coastlines and weaker winds farther north. Its signature in the mean sea ice drift patterns, with larger sea ice velocities along the Alaskan coastline, and in the SSH field, with strong SSH gradients is more confined to the Canada Basin in the east-west direction (Figures 4b and 4c). This is particularly true along the Northwind Ridge when compared to the Arctic atmospheric high and to a lesser extent to the ice circulation, thus suggesting the possible important contributions of the thinner sea ice cover (which facilitates air-sea momentum fluxes) along the Alaskan coastline and of the bathymetry-constrained surface currents in the ice momentum balance.

## 4.1. Climatological Circulation in the Beaufort Gyre

The circulation of the sPW layer shows a closed anticyclonic circulation, similar to that of the surface gyre seen in SSH with a smaller spatial extent (Figure 4d, see Figure S3a in Supporting Information S1 for the magnitude of the derived geostrophic velocities). Adjacent to a closed recirculation in the centre of the Canada Basin is a flow along streamlines extending from the Chukchi shelf to the Canadian Arctic Archipelago (CAA), much of which flow merges with the Transpolar Drift (a surface current transporting sea ice from the Eurasian shelf to Fram Strait). This flow is joined by waters of Eurasian origin, which flow eastward along the Siberian shelf before turning poleward, where they either exit the layer through an outcrop or join the Pacific Waters and continue toward the CAA. Consistent with this is work by Alkire et al. (2007) who show a front separating these two water masses between Mendeleev and Alpha Ridges and Lomonosov Ridge, although the confluence of Chukchi and Eurasian Waters is found further to the south in our analysis. This front was documented in previous studies and its location was shown to change with the main atmospheric patterns. A southern position for the 2005–2017 decade is coherent with, for example, Steele et al. (2004) and a general positive Arctic Oscillation(AO) for these years. This general picture is robust to the choice of SSH product used in defining the Montgomery potential. The choice of hydrography, however, has a larger influence on outcrop locations and, because of this, on Eurasian water pathways on the sPW layer (Figure 6).

A closer look at the Canada Basin (Figure 5a) shows that in contrast to SSH, contours of $M$ are seen to intersect the Alaskan shelfbreak. Near Pt. Barrow, for example, the time mean flow appears to stem from the Alaskan and Chukchi shelves, consistent with the possibility of waters ventilating the gyre at this location (e.g., Timmermans et al., 2017). This picture is robust to the choice of hydrography but not to the choice of SSH. As mentioned, O-SSH presents a stronger gradient along Chukchi shelfbreak and shows a stronger recirculation which, here, translates into a stronger westward flow (with time mean $M$ contours aligned with the Alaskan coast; Figure 6). Also shown in Figure 5a are seasonal outcrops. From these, it is evident that only a relatively small fraction of the recirculation remains below the mixed layer year-round. Put another way, most of the sPW portion of the Beaufort
<figure><img src="image_6.png" /><figcaption>**Figure 4.** (a) Mean sea level pressure from ERA-interim. (b) Mean sea ice velocity direction (arrows) and normalized speed (color shading) from the Polarpath Finder (Tschudi et al., 2020). (c) Mean W-SSH anomaly times the acceleration of gravity. Montgomery potentials (M, contours) and conservative temperature ($T_C$, color shading) for the sPW layer (d), wPW layer (e), and AW layer (f) from WOA and W-SSH anomaly. For reference, the magnitude of the geostrophic velocities within the sPW layer along Chukchi shelfbreak is 5 cm/s, and in the centre of the Canada Basin is 1 cm/s, see Figure S3 in Supporting Information S1. All data sets are taken for the time period 2005–2017. Dashed black lines indicate the outcrop of the isopycnal layer into the mixed layer, defined with a potential density difference $\Delta\sigma_0 = 0.05\text{kg/m}^3$ from the surface (Timmermans et al., 2017). Gray plain lines correspond to 100 and 1,000 m isobaths. Note the different color scales of M, $T_C$ associated with each layer.</figcaption></figure>
Gyre is seen to either be ventilated at outcrops along the Chukchi shelf or outcrops in the mixed layer during certain months. Temperatures within the sPW are relatively high throughout the Canada Basin and a particular "hot spot" of warm water is found east of the Northwind Ridge. That the temperature field is not generally aligned with *M* contours implies that ventilation pathways must also include processes operating on time and spatial scales not resolved here. We address seasonal variability in the large scale circulation in Section 4.2, but inter-annual variability and small-scale eddies and filaments may also likely play a role as highlighted by, for example, recent in situ measurements of surface temperature showing filaments and eddies of warm sPW breaking off at Point Barrow (e.g., Kubryakov et al., 2021; MacKinnon et al., 2021) or pieces of evidence of year to year variability in the circulation, such as the recent acceleration of the Beaufort Gyre around 2007–2010 (Armitage et al., 2018; Regan et al., 2019).

Finally, we note that a well known feature of sPW—the narrow (~15 km) and intermittent tongue of warm water reported flowing east along the Alaskan coastline in Steele et al. (2004) and Brugler et al. (2014) and forming the Alaskan Coastal Current—is not present in our analysis despite its density corresponding to that of the sPW layer. This is not surprising given that our spatial resolution is only 1/4° and that intermittent features may or may not appear in the climatological average. We schematically represent the main climatological geostrophic pathways highlighted from the Montgomery potential for the 2005–2017 decade in the sPW layer on Figure 7a.

Circulation in the wPW layer is more sensitive to the choice of data sets. Using the WOA and W-SSH data sets, the centre of the gyre does not show a robust closed anticyclonic circulation; instead, the time mean flow shows more of a stagnant pool (or broken into two smaller recirculations, Figure 5b). Around this pool is a circulation extending from the Alaskan and Chukchi Shelves, the inner part of which reconnects with the Alaskan Shelf (presumably joining an unresolved boundary current there), with the outer part flowing into the CAA. A schematic representation of these main pathways is visible on Figure 7b. Although these key features are robust to changes in the SSH data set and hydrography, the strength of the recirculation, however, is sensitive to the choice of hydrography. MIMOC has a more pronounced fresh water “bowl” and therefore produces stronger recirculation. Similarly, subduction suggested by contours of *M* intersecting the Alaskan and Chukchi Shelves is not diagnosed when using O-SSH, which instead shows a resolved boundary current flowing anticyclonically along the continental margin (Figure 6).

As was the case for sPW, waters flowing from the Chukchi Shelf toward the CAA are joined by waters of Eurasian origin (Figure 4e, see Figure S3b in Supporting Information S1 for the magnitude of the derived geostrophic field). These are colder and shallower than the Chukchi-origin waters (not shown), suggesting a front along the northern side of Mendeleev and Alpha Ridges. In the Nansen and Makarov Basins, the Eurasian Waters flow cyclonically along the bathymetry with, similarly to the sPW, some of the flow outcropping in Nansen Basin and some joining the Chukchi-origin waters in a current flowing from the Laptev Sea to the CAA. These waters of Eurasian origin outflow primarily through Fram, Nares Straits and Queen Elizabeth Islands Straits, whereas the Chukchi-origin waters appear to veer westward upon reaching the Archipelago thus suggesting outflows through the M'Clure, Amundsen, and Queen Elizabeth Island Straits, in accord with simulation results (e.g., Hu & Myers, 2013; Lique et al., 2010). Alternatively, a boundary current along the Archipelago could recirculate part of these waters without being resolved in our description of the circulation as indicated by streamlines connecting with the shelfbreak. These outflows are robust to changes in SSH data sets and hydrography (Figure 6), and generally agree with the nutrient signature measured at Fram and Nares Strait (Forryan et al., 2019; Jones et al., 2003). However, previous studies reported an outflow of Pacific origin at Fram Strait (Falck, 2001; Falck et al., 2005), whereas our results suggest the outflow there to be mainly, on decadal average, of Eurasian origin. Whether this difference follows from the very similar nutrient signatures of shelf water masses from Bering and Laptev Sea, or from the large interannual variability in the Pacific Waters export that is not captured by our representation of the circulation (Alkire et al., 2007; Jahn et al., 2010) is left for future analysis.

In the Canada Basin, and irrespective of the data sets used, our reconstruction shows only a weak time mean circulation in the AW layer (Figures 4f, 5c, and 6, Figure S3c in Supporting Information S1). Using the MIMOC hydrography, a large stagnant pool of AW is found to lie beneath the surface and PW gyre. This is flanked by an anticyclonic flow of water that, following a meandering path, can be traced back to the Nansen and Amundsen Basins. A similar flow of AW from the Eurasian to Canada Basins is also seen (but is weaker) when using WOA. A similar pattern is visible in the late 20th century circulation as depicted by, for example, Karcher et al. (2012), although in our reconstruction it penetrates into the Canada Basin. The Nordic Seas show a well-defined current
<figure><img src="image_7.png" /><figcaption>**Figure 5.** Same as Figures 4d–4f but zoomed on the Canada Basin. Dashed black lines represent the most restrictive limit of subduction (outcrops) over the 12 months. Dashed black lines represent the outcrops of the mean annual WOA hydrography.</figcaption></figure>

along the bathymetry and a consistent inflow of AW from the shelf at St. Anna Trough. In the Nansen and Amundsen Basins, a less well-defined cyclonic circulation is evident, consistent with a flow following bathymetric features, as described by Isachsen et al. (2003). Over and to the west of the Northwind Ridge is some evidence for an anticyclonic circulation around the ridge (with a mirroring, but weaker cyclonic circulation on the eastern side). Further east, along the Alaskan continental margin, there is evidence of a cyclonic (eastward) current banked up against the shelf, consistent with previous results from, for example, Karcher et al. (2012) and Woodgate et al. (2007). Note, however, that this feature disappears when O-SSH is used in place of W-SSH. To summarize, results for the AW layer (especially in the Canada Basin) are not robust. Even the sign of the overall circulation (cyclonic vs. anticyclonic) is unclear, and Figure 7c highlights the main pathways using WOA and W-SSH only.

## 4.2. Seasonal Variability and Ventilation From the Alaskan Shelf

The data are insufficient to represent interannual variability, but do allow us to attempt describing a seasonal cycle. We focus on the Canada Basin and begin with the sPW, presenting monthly climatologies in Figure 8. In some instances, sparseness of the data appears to have introduced spurious features, for example, the strong local minimum of *M* in March is seen just north of the Yukon, as shown in Figure 8c. The strong seasonality in shape, extent, and location of the Beaufort Gyre seen in the figure results from a strong seasonality in near-surface salinity (not shown). That is, seasonality in SSH does not account for the strong seasonality we see in the Montgomery potentials (Figure S1 in Supporting Information S1; see also Figure 3 from Regan et al. (2019)).

The anticyclonic circulation described in Section 4.1 for the sPW layer is also evident in seasonal averages (Figure 8), albeit with changes in location, shape, and strength (see below). The wPW decadal circulation can be thought of as an average of localized (and sometimes, strong) patterns that vary across the seasons (Figure 9).

A closed circulation within the basin is well defined from June to January, and then opens up and deforms between February and May. The closed circulation is confined to the Canada Basin in summer and spills over Northwind Ridge in winter. Streamlines emanating from the Alaskan Shelf tend to flow roughly parallel to the continental margin from November to May, and then turn to flow northward toward the basin interior in summer months. This northward shift in the direction of streamlines during summer is consistent with warm waters ventilating the gyre interior after subducting into the sPW layer near Pt. Barrow, as suggested by Timmermans et al. (2017). Additional ventilation is found in late spring/summer at outcrops to the west/north of Northwind Ridge. The outcrops there are co-located with positive anomalies of sea surface salinity (not shown); the cause of these salinity anomalies is yet to be determined.

The wPW also shows strong seasonal variability (Figure 9), some of which seems to align well with that found in the sPW layer. Montgomery potentials consistent with subduction along the Alaskan coastline and Chukchi shelf-break into the Canada basin, for example, are found between May and

September, when the upper sPW Montgomery potential contours tilt northward, so that geostrophic flow is toward the basin interior. Winter months also show penetration of wPW in the CB basin along the shelf break (January–March), in line with the seasonal subduction of sPW and wPW according to the seasonal migration of
<figure><img src="image_8.png" /><figcaption>**Figure 6.** Montgomery potential (M, contours) and conservative temperature (Tc, color shading) for the sPW layer (first row), wPW layer (second row), and AW layer (third row) from the combined MIMOC/W-SSH (first column), WOA/W-SSH (second column), and WOA/O-SSH (third column) data sets for the time period (2005–2017). Dashed black lines indicate the outcrop of the isopycnal layer into the mixed layer, defined with a potential density difference $\Delta\sigma_0$ = 0.05kg/m$^3$ from the surface. Gray lines indicate the 100 and 1,000 m isobaths. Note the different color-scales associated with each layer.</figcaption></figure>

outcropping isopycnals characterizing these water masses as documented by Timmermans et al. (2017). Another example is along the western flank of the Northwind Ridge, where an intense anticyclonic recirculation is seen between March and July. This is roughly aligned with outcrops of the sPW layer, and with a similar anticyclonic circulation in June and July in the AW layer (not shown). Not having an obvious counterpart in the sPW layer, the wPW inflow from Pt. Barrow is seen to strengthen and veer to the east, forming an anticyclonic current along the shelfbreak. A Lagrangian perspective of the wPW is visible on Figure S2 in Supporting Information S1.
<figure><img src="image_9.png" /></figure>

<figure><img src="image_10.png" /></figure>

<figure><img src="image_11.png" /><figcaption>Figure 7.</figcaption></figure>

Of particular interest to us is how warm summer waters from Pt. Barrow and other locations along the Alaskan shelf ventilate the sPW layer. The summertime subduction of sPW along the Chukchi and Alaskan shelves, together with large springtime outcrops along Northwind Ridge suggest geostrophic pathways for heat to be transported toward the center of the Canadian Basin. Such subduction was documented along the shelf break from in situ vertical hydrographic profiles (Timmermans et al., 2014, 2017, 2018). With this in mind, we consider particle trajectories calculated assuming flow along the sPW density surface, and using geostrophic velocities inferred from the Montgomery potentials. Specifically, we interpolate monthly velocities to obtain semi-monthly fields and use these with a second order Crank Nicholson advection scheme to calculate Lagrangian trajectories. Particles are initialized for each of the summer months along the Alaskan and Chukchi shelves, and are advected for 3 years or until they reach an outcrop. Figures 10a–10d shows the trajectories to move anticyclonically around the basin, as expected, but they never enter the basin interior (i.e., the region of closed time-mean *M* contours). Instead, most trajectories are seen to outcrop near the Northwind Ridge. Initializing the particles downstream of the outcrop area gave similar results with anticyclonic paths that outcrop upon return in the same area (not shown).

This is perhaps unsurprising since many other advective processes are also at play. These include mesoscale eddies and filaments such as those recently observed near Pt. Barrow (MacKinnon et al., 2021), along the Alaskan coastline (Kubryakov et al., 2021), or in idealized models of the baroclinically unstable Alaskan shelfbreak (Spall et al., 2018). These, of course, are not captured by the WOA climatology. Additionally, large-scale interannual variability is likely comparable in magnitude to the seasonal variability discussed above. Not only is there significant variability in the location, shape, and strength of the gyre (Armitage et al., 2018; Regan et al., 2019) but also in the wind direction and strength along the Alaskan coastline (westerlies vs. easterlies) that controls the extent and strength of the Alaskan Coastal Current (Brugler et al., 2014). Finally, the area where most of the particles were found to outcrop above is close to the Northwind Ridge, where there is known to be a weak but persistent Ekman downwelling (albeit with some year-to-year variability; Meneghello et al., 2018).

With these processes in mind, we recalculate the trajectories adding noise to the velocity field. The noise is intended as a crude model of some of the components of the flow mentioned above that are unresolved in our analysis. Specifically, at each (semi-monthly) time step and each point in space we add a random perturbation to the velocity field used to advect the trajectories. The perturbation has an amplitude chosen randomly from a uniform distribution between [0, *u*\*] with *u*\* a chosen constant velocity, and a direction oriented randomly in the plane tangent to the sPW isopycnal surface. For each initial position and constant value of *u*\*, we calculate 50 trajectories. Results

**Figure 7.** Schematic of the geostrophic circulation for the period 2005–2017 on three isopycnal layers corresponding to the core of the summer Pacific Waters (sPW), the winter Pacific Waters (wPW) and the Atlantic Waters (AW), diagnosed from Montgomery potentials using the combined WOA/W-SSH. Blue stands for anticyclonic circulation, red stands for cyclonic circulation. Dashed black lines indicate the outcrop of the isopycnal layer into the mixed layer, defined with a potential density difference $\Delta\sigma_0$ = 0.05kg/m$^3$ from the surface. Gray lines indicate the 100 and 1,000 m isobaths.
<figure><img src="image_12.png" /><figcaption>**Figure 8.** Seasonality of Montgomery potential (*M*, contours) and conservative temperature (*T*<sub>C</sub>, color shading) for the sPW layer in the combined WOA/W-SSH data set for the period 2005–2017 (zoomed on the Canada Basin). Dotted black lines indicate the outcrop of the top isopycnal layer into the mixed layer, defined with a potential density difference Δ*σ*<sub>0</sub> = 0.05kg/m<sup>3</sup> from the surface. Gray lines indicate the 100 and 1,000 m isobaths.</figcaption></figure>
<figure><img src="image_13.png" /><figcaption>**Figure 9.** Seasonality in Montgomery potential (M, contours) and conservative temperature (Tc, color shading) for the wPW layer in the combined WOA/W-SSH data set 2005–2017 (zoomed on the Canada Basin). Dashed black lines indicate the outcrop of the isopycnal layer into the mixed layer, defined with a potential density difference Δσ₀ = 0.05kg/m³ from the surface. Gray lines indicate the 100 and 1,000 m isobaths.</figcaption></figure>
<figure><img src="image_14.png" /><figcaption>**Figure 10.** 3-year Lagrangian tracks initialized in the sPW layer at Point Barrow (different colors, along longitude) for all summer month advected using the geostrophic velocities without (first row) and with (second row) a random perturbation in the velocities magnitude and direction ($u^* = 1.5$ cm/s); and for initialization in September only with $u^* = 0.5, 1.0, 2.0$, and $3.0$ cm/s (third row). Initialization location is indicated with a black circle, with second and third rows showing 50 initializations of the same particles.</figcaption></figure>

(Figures 10e–10l) show that the noise allows some of the particle trajectories to avoid the outcrop and continue around the gyre, with this effect being more pronounced for particles released in September or October than in July or August. It is also noteworthy that a relatively large amplitude ($>1.15–2$ cm/s) for $u^*$ is needed before the particles were seen to enter into the centre of the gyre.

Although this idealized perturbed velocity is only a toy experiment that does not represent diapycnal processes, nor vertical transient fluxes and thus should be analyzed with care, it demonstrates the need for across streamlines transport to feed the Canada Basin interior with warm shelf waters, as this is not sustained by the geostrophic along-isopycnal flow. A precise quantification of this transport as well as disentangling mesoscale from inter-annual variability is beyond the scope of this paper - and beyond the capacities of this data set and is thus left for future analysis.

**5. Conclusion**

Despite their importance for the local heat balance and the sea ice cover above, the pathways of the relatively warm summer Pacific Waters (sPW) and Atlantic Waters (AW) that feed the subsurface heat reservoirs in the Canada Basin, and those of the cold winter Pacific Waters (wPW) that build the halocline in the Amerasian basin of the Arctic Ocean remain poorly documented from observations. To this end, we use available hydrographic and sea surface height observations to calculate the Montgomery potential ($M$) on isopycnal surfaces representative of these three water masses. Contours of $M$ correspond to streamlines of the geostrophic flow which are used to reconstruct estimates of the time mean circulation and its seasonal variability for the period 2005–2017.

The basin-scale geostrophic circulation shows the strong and robust anticyclonic circulation of the surface waters in the Canada Basin to rapidly decay with depth. In the sPW, it remains clearly visible although only a small fraction of the basin is characterized by closed streamlines which remain below the surface mixed layer. Over most of the gyre, flow is instead ventilated either from the Alaskan and Chukchi Shelves or from outcrops in the
basin interior, near Northwind Ridge. In the wPW layer, the gyre weakens considerably, to the extent that in the core of the gyre, variability dominates over the mean circulation. Also evident (although not robust to the choice of data sets used) is a cyclonic boundary current in the AW layer that follows the Alaskan coast and an anti-cyclonic circulation around Northwind Ridge. These results are summarized in Figure 10.

A robust signal suggestive of subduction and ventilation of the sPW and wPW from near Point Barrow and along the shelfbreak in the Chukchi Sea is evident in the time mean fields. This is robust to the choice of hydrography (WOA or MIMOC) but not to the choice of SSH product (it is only visible when using the satellite SSH data set of Armitage et al. (2016) but not the model-derived ORAS5 reanalysis). Contours of *M* stemming from the continental margin are mostly aligned with the shelfbreak except in summer when they tilt toward the interior of the Canada Basin. Tracing particle trajectories are released from this area in summer; however, they do not show them bringing warm water into the core of the gyre. Instead, most trajectories were found to outcrop near the Northwind Ridge. Adding noise to the resolved time mean and seasonal velocities to mimic unresolved processes does allow for some particles to avoid the outcrop and, for sufficiently strong values of the noise, to enter into the core of the gyre—thus pinpointing the possible key role of these unresolved processes in sPW ventilation and heat subduction.

The coarse temporal resolution of both data sets (decadal) precludes any robust quantification of the interannual variability of the geostrophic flow. As a consequence, the possible anticyclonic intensification over 2007–2011 documented from satellite observations (e.g., Regan et al., 2019) is not visible in our analysis. It is also likely that our climatology spatially averages important variability in the time mean circulation. For example, when considering the AW circulation in the central Arctic, possible incursions above Lomonosov and Mendeleev Ridges at different locations (e.g., Karcher et al., 2012) are likely unresolved.

In sum, we offer a decadal-scale average with a crude representation of the seasonal cycle of the upper layer flows within the Amerasian Basin, relying on observations. Interannual variability of these pathways, despite being suggested as important in the literature, is beyond the scope of this paper and is left for future work. Similarly, a more comprehensive quantification of the heat pathways that fill the sPW and AW heat reservoirs and cold wPW pathways that build the halocline between the two, as well as disentangling the relative roles of mesoscale eddies and filaments and interannual variability are also left for future analysis.

## Data Availability Statement

The World Ocean Atlas (WOA) 2018 climatology was downloaded from Boyer et al. (2018). The Monthly Isopycnal & Mixed-layer Ocean Climatology (MIMOC) was downloaded from Schmidtko et al. (2013). Arctic dynamic topography/geostrophic currents data were provided by the Centre for Polar Observation and Modelling, University College London (Armitage et al., 2016, 2017). Ocean Reanalysis System 5 Sea Surface Height data set was downloaded from Zuo et al. (2019). The scripts used to construct the hybrid SSH field and the Montgomery potentials on isopycnal coordinates are available at Planat (2024).

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## Acknowledgments

NP was supported by the Fonds de recherche du Québec—Nature et technologie (FRQNT) through a Doctoral Training Scholarship, a Natural Sciences and Engineering Research Council of Canada (NSERC) Accelerator Supplements (Grant RGPAS/2018-522502 and the Canada Research Chair program awarded to COD, and an NSERC Discovery Grant awarded to BT (Grant RGPIN/2018-04838). DS was supported by an NSERC Discovery Grant (Grant RGPIN/2022-03401). NP received a stipend from Québec-Océan and we gratefully acknowledge scientific and financial support from Québec-Océan. We thank two anonymous reviewers for their constructive comments on the manuscript.
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