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Methane and nitrous oxide distributions across the North American Arctic Ocean during summer, 2015 Fenwick, Lindsay Alexandra 2016

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   Methane and nitrous oxide distributions across the North American Arctic Ocean during summer, 2015  by Lindsay Alexandra Fenwick B.Sc. (Hons) Queen’s University, 2013 A THESIS SUBMITTED IN PARTIAL FULFILLMENT  OF THE REQUIREMENTS FOR THE DEGREE OF   Master of Science in  The Faculty of Graduate and Postdoctoral Studies  (Oceanography)  The University of British Columbia (Vancouver) July 2016 © Lindsay Alexandra Fenwick, 2016     ii Abstract We collected Arctic Ocean water column samples for methane (CH4) and nitrous oxide (N2O) analysis on three separate cruises in the summer and fall of 2015, covering a ~10,000 km transect from the Bering Sea to Baffin Bay. Our sampling program provides a large-scale, three-dimensional view of methane and nitrous oxide concentrations across contrasting hydrographic environments, from the deep oligotrophic waters of the deep Canada Basin and Baffin Bay, to the productive continental shelf regions of the Bering and Chukchi Shelves.  Percent saturation relative to atmospheric equilibrium ranged from 30-800% and 75-145% for CH4 and N2O, respectively, with the highest concentrations of both gases occurring in waters overlying the continental shelf in the northern Chukchi Sea. Nitrification and denitrification in the sediments of the Bering and Chukchi Shelves constituted a major source of N2O to the water column, and the resulting high N2O concentrations were transported across the entire North American Arctic. Methane sources to the water column were more spatially heterogeneous, reflecting a greater variety of hydrographic influences, including likely inputs from sediments, rivers, and sea ice processes. Localized regions of high CH4 concentrations were observed at various locations across our sampling transect, but unlike N2O, CH4 was rapidly consumed through microbial oxidation, as shown by the 13C enrichment of CH4 at low concentrations.  High CH4 signatures were thus more localized across our sampling region.  For both CH4 and N2O, surface super-saturation and sea-air fluxes were generally low across the region, with valuess of 1.3 ± 1.2 µmol m-2 d-1 and -0.52 ± 1.0 µmol m-2 d-1, for CH4 and N2O, respectively.  Low surface water concentrations were at least partially attributable to dilution by low CH4 and N2O fresh water.  Our results provide insight into the factors controlling the distribution of CH4 and N2O in the North American Arctic Ocean, and an important baseline data set against which future changes can be assessed.    iii Preface This dissertation is based on a concept formed by my supervisor (Dr. Philippe Tortell) and myself. Fieldwork and sample collection was done by me, Dr. Tortell, and Department of Fisheries and Oceans scientists, Sarah Zimmermann and Bill Williams. Nutrient samples and CTD data for SWL 2015-09 and LSSL 2015-09 were collected and analyzed by the Institute of Ocean Sciences. Ammonium samples for SWL 2015-09 were collected by Lee Cooper and analyzed by the Nutrient Analytical Services Laboratory at the University of Maryland. Nutrient samples for ArcticNet 1502 were collected and analyzed by the research group of Jean Eric Tremblay (Université Laval). Ellen Damm (Alfred Wegner Institute, Bremerhaven, Germany) analyzed the ∂13CCH4 samples and helped interpret the data. I was responsible for the analysis of CH4 and N2O samples using GC-MS, interpretation of data, and writing of the dissertation. Dr. Tortell thoroughly edited the thesis, provided advice and ideas, and helped interpret data. David Capelle, Ellen Damm, Kristina Brown, and Svein Vagle contributed valuable edits. A version of this thesis will be submitted as a manuscript with co-authors Philippe Tortell, David Capelle, Ellen Damm, Svein Vagle, Bill Williams, and Sarah Zimmermann.     iv Table of Contents Abstract ............................................................................................................................................ ii	  Preface ............................................................................................................................................. iii	  Table of Contents .......................................................................................................................... iv	  List of Tables ................................................................................................................................... v	  List of Figures ................................................................................................................................. vi	  List of Abbreviations ................................................................................................................... viii	  Acknowledgements ....................................................................................................................... ix	  Dedication ....................................................................................................................................... x	  1   INTRODUCTION ................................................................................................................................ 1	  1.1	   CH4 AND N2O IN THE OCEAN ..................................................................................... 1	  1.2	   CH4 AND N2O IN THE ARCTIC OCEAN ........................................................................ 4	  2   METHODS AND REGIONAL SETTING ........................................................................................... 8	  2.1	   FIELD SAMPLING ......................................................................................................... 8	  2.2	   DISSOLVED GAS CONCENTRATIONS AND δ13CCH4 VALUES ........................................ 8	  2.3	   NUTRIENT ANALYSIS .................................................................................................. 9	  2.4	   BIOGEOCHEMICAL CALCULATIONS ......................................................................... 10	  2.5	   SEA-AIR FLUX CALCULATIONS ................................................................................. 10	  2.6	   REGIONAL SETTING .................................................................................................. 11	  2.6.1	   Bering and Chukchi Seas ............................................................................. 11	  2.6.2	   Canada Basin and Beaufort Sea .................................................................. 13	  2.6.3	   Canadian Arctic Archipelago ..................................................................... 14	  2.6.4	   Seafloor methane sources ............................................................................ 14	  3   RESULTS AND DISCUSSION .......................................................................................................... 16	  3.1	   NITROUS OXIDE ........................................................................................................ 16	  3.1.1	   N2O in the Bering and Chukchi Seas ......................................................... 18	  3.1.2	   N2O in the Canada Basin and Beaufort Sea .............................................. 23	  3.1.3	   N2O in the Canadian Arctic Archipelago ................................................. 26	  3.2	   METHANE ................................................................................................................. 27	  3.2.1	   CH4 in the Bering and Chukchi Seas ......................................................... 28	  3.2.2	   CH4 in the Canada Basin and Beaufort Sea .............................................. 33	  3.2.3	   CH4 in the Canadian Arctic Archipelago .................................................. 36	  3.3	   SEA-AIR CH4 AND N2O FLUXES .................................................................................. 38	  4   CONCLUSIONS ................................................................................................................................. 42	  4.1	   MAIN FINDINGS ........................................................................................................ 42	  4.2	   CAVEATS AND LIMITATIONS .................................................................................... 43	  4.3	   FUTURE OUTLOOK ................................................................................................... 44	  References ...................................................................................................................................... 46	   v List of Tables Table 1:	   Cruise details ............................................................................................................ 8	  Table 2:	   Sea-air flux of N2O ................................................................................................ 39	  Table 3:	   Sea-air flux of CH4 ................................................................................................. 40	  Table 4:	   Regional annual fluxes of CH4 and N2O ............................................................ 41    vi List of Figures Figure 1: Map of study area, showing the locations of depth profiles (yellow circles) and dominant Pacific Water boundary currents. The black line denotes the transect along which the contour plots (Figure 5 and Figure 10) were derived. Based on [Macdonald et al., 2004]. ............................................................................................................................ 7	  Figure 2: Salinity vs. temperature plot showing the different water masses observed in the North American Arctic. The top panel shows percent saturation of N2O, while the bottom panel shows percent saturation of CH4. Circles, squares, and triangles denote measurements from the Bering and Chukchi Seas, Canada Basin, and Canadian Arctic Archipelago, respectively. ................................................................................................... 17	  Figure 3: Density vs. concentration plots averaged by sub-region. The grey shaded area is the standard deviation of the mean. ................................................................................. 18	  Figure 4: Map and corresponding N2O depth profiles in the Bering and Chukchi Seas. Black filled circles denote a measurement, while the error bars represent the standard deviation of the duplicate samples. The grey dotted line is the atmospheric equilibrium concentration. ................................................................................................ 20	  Figure 5: Nitrous oxide and density contour plots. The top two panels show the distribution of N2O and density along a ~1500km transect from the Bering to the Beaufort Sea (black line on Fig. 1), respectively. The bottom panel shows the mean concentration of N2O on the 26.5 density surface, illustrating the accumulation of N2O along this isopycnal. .............................................................................................................................. 21	  Figure 6: ∆N2O vs. AOU and N* plots, broken down by sub-region. Only the Canada Basin UHL is shown in order to better compare the properties of the Pacific water masses between regions. Trend lines are shown where there is a statistically significant regression model (p<0.05). ................................................................................................ 23	  Figure 7: Map and N2O depth profiles in the Canada Basin and Beaufort Sea. .................. 24	  Figure 8: Map and depth profiles of N2O in the CAA. ........................................................... 27	  Figure 9: Map and depth profiles of CH4 in the Bering and Chukchi Seas .......................... 29	   vii Figure 10: Methane and density contour plots. The top panel shows the distribution of CH4 along a ~1500 km transect from the Bering to the Beaufort Sea (black line on Fig. 1). The bottom panel shows the density distribution along the same transect. ............... 30	  Figure 11: Oxidation curve calculated using a Rayleigh distillation model with an initial concentration of 31 nmol/L, a ∂13CCH4 value of -40‰ VPDB, and an isotopic fractionation factor of 1.002. .............................................................................................. 32	  Figure 12: Map and depth profiles of CH4 concentrations in the Canada Basin and Beaufort Sea .......................................................................................................................................... 33	  Figure 13: This is a representative depth profile of CH4 and fluorescence, showing the co-occurrence of the CH4 and chlorophyll maxima. Filled black circles denote CH4 concentration and error bars represent the standard deviation of the duplicates. The dotted line is fluorescence. ................................................................................................. 34	  Figure 14: Map and depth profiles of CH4 concentrations in the CAA and Baffin Bay ..... 37	  Figure 15: CH4 vs. salinity scatter plot of all samples collected in the CAA and Baffin Bay ................................................................................................................................................ 38	     viii List of Abbreviations ACW Alaska Coastal Water AOU Apparent oxygen utilization AW Atlantic Water BSW Bering Sea Water CAA Canadian Arctic Archipelago CH4 Methane N2O Nitrous oxide NH4+ Ammonium NO3- Nitrate NO2- Nitrite O2 Oxygen PML Polar Mixed Layer PSW Pacific Summer Water PW Pacific Water PWW Pacific Winter Water UHL Upper Halocline Layer      ix Acknowledgements I have been fortunate to be supported by many wonderful individuals without whom this thesis would not have been possible. First and foremost, I would like to thank my supervisor, Philippe Tortell, for his unwavering encouragement, enthusiasm, and ability to thoroughly edit a 10,000-word manuscript sent at midnight by 6am the next morning. Thank you for the initiation into the world of oceanography. I would also like to thank David Capelle for his ridiculous accents, incredible patience, and calm demeanor while being informed that the instrument is no longer working, yet again. No number of washed bottles will ever begin to repay you. I am also very appreciative of the insightful editing by Ellen Damm and Kristina Brown, without which this thesis would not be nearly as strong. Thank you to Kristin Orians and Roger François for their revisions of this thesis. The dissolved gases in this thesis would still be in the ocean had it not been for the captains and crews of the CCGS Sir Wilfrid Laurier, CCGS Louis St. Laurent, CCGS Amundsen, and CCGS John P. Tully, and the scientists at IOS. Thank you in particular to Dr. Svein Vagle, for his can-do attitude and enthusiasm for this project, and Sarah Zimmermann and Bill Williams for their dedication and insights into Arctic hydrography. Thank you to the Grebmeier group, Jennifer Long, and Curtis Martin for their camaraderie in the Arctic and musical puppet show talents, and Bruce Patterson, for helping me protect my egg. I am indebted for the friendship, insight, and caring of the Tornado lab, my officemates, and many others at UBC. Lian, Kang, Joanne, Dave, Nina, Tereza, Ania, Anna, Alysia, Robert, Will, Manuel, Lucy, Michael, Brett, and Kirsten- I will miss you all.  None of this would have been possible without my family. My parents have always been supportive of my every interest and encouraged me to collect data- of any kind. My grandparents have been a source of inspiration, and their relentless support of education has lead to my achievement of the same lever of education as my Grandpa.  And D.J., my best friend and teammate- I learn from you every day. I could not have done this without your love and emotional support. Thank you.     x Dedication    for Elaine and Don 1 1    Introduction After carbon dioxide, methane (CH4) and nitrous oxide (N2O) are the two greenhouse gases that have most influenced climate warming over the past 250 years [IPCC, 2013]. While atmospheric concentrations of CH4 and N2O have been relatively stable for the past 12,000 years, transitions from glacial to interglacial climates have previously been accompanied by large atmospheric increases in both of these gases. More recently, the atmospheric concentrations of these gases have increased by 165% and 20%, respectively, since preindustrial times, as a result of various anthropogenic activities. The current values (1910 and 328 ppb for CH4 and N2O, respectively) have no precedent in the past 800,000 years (Data provided by NOAA ESRL Global Monitoring Division, Boulder, Colorado, USA; http://esrl.noaa.gov/gmd/) [Schilt et al., 2010]. Although the contribution of CH4 and N2O to radiative forcing in the atmosphere is clear, the cycling of these gases among various global reservoirs  is still relatively poorly understood. Understanding the global budget and cycling of CH4 and N2O is a primary research objective in the Earth Sciences. 1.1   CH4 AND N2O IN THE OCEAN The ocean plays a key role in the global cycles of greenhouse gases, including N2O, and to a lesser extent, CH4. Surface ocean waters are estimated to account for ~2 and 20-60 % of the total emissions of CH4 and N2O to the atmosphere, respectively, resulting from a series of complex production and consumption processes [Reeburgh, 2007; IPCC, 2013]. A variety of biologically-mediated processes produce and consume CH4 and N2O within the ocean and the sediments that lie underneath. Depending on environmental conditions, the ocean can be a net source or sink of these gases.  2 N2O cycling in the marine environment is controlled by two processes: nitrification and denitrification. The former, which is thought to be the dominant source of oceanic N2O, is completed in two steps. The first, carried out by ammonia-oxidizing bacteria, oxidizes ammonium (NH4+) to NO2-. The decomposition of NH2OH, an intermediate in the oxidation of NH3 to NO2-, produces N2O as a by-product. The second step, performed by nitrite-oxidizing bacteria, oxidizes NO2- to NO3-. Although nitrification is an aerobic process, N2O yields are enhanced under low oxygen conditions. This is thought to be because of the “nitrifier-denitrification” process, or the tendency of nitrifiers to reduce NO2- to N2O when they are O2-limited [Codispoti et al., 2005]. Denitrification is another key process in the N2O cycle. This pathway is limited to suboxic to anoxic conditions, and reduces fixed nitrogen (NO3-) to nitrogen gas (N2), producing N2O as an intermediate. At oxygen concentrations near the upper limit for denitrification, N2O is mostly released, resulting in a net production of N2O [Codispoti et al., 2005]. Under low oxygen conditions, however, well-established denitrification systems proceed to completion, resulting in the biological consumption of N2O. The highest production of N2O occurs at the boundaries of oxygen minimum zones, where nitrification and incomplete denitrification occur simultaneously, yielding elevated N2O concentrations [Codispoti et al., 2005]. Sources of CH4 to the ocean are varied, including inputs from gas hydrates (frozen CH4 structures), abiotic serpentinization at mid-ocean ridges, and seepage from hydrocarbon deposits [Reeburgh, 2007]. The majority of CH4 in the ocean likely is produced though sedimentary methanogenesis, during the anaerobic breakdown of organic matter [Reeburgh, 2007], and subsequent diffusion into the water column. Additional sources of CH4 may come from freshwater runoff from regions with anaerobic sediments. Despite relatively large contributions from these sources, the CH4 concentrations in much of the ocean remain below atmospheric saturation due to intense aerobic and anaerobic microbial oxidation in both the sediments and the water column. However, somewhat paradoxically, the well-oxygenated surface layer of the ocean sometimes contains a CH4 maximum. Several hypotheses have been proposed to explain this enigma, including CH4 production in the anoxic microenvironments such as the guts of zooplankton or within sinking particles [eg., Oremland, 1979; De Angelis and Lee, 1994; Karl and Tilbrook, 1994], the cleavage of methylated substrates by bacteria or phytoplankton [Karl et al., 2008;  3 Damm et al., 2010], and advection from regions with hydrocarbon seeps or destabilized gas hydrates [Naqvi et al., 2010 and references therein]. The cause(s) of this apparent 'methane paradox' remain poorly understood. Kinetic isotope effects resulting from the production and oxidation of CH4 impart a signature on the stable isotopic composition of CH4 which can be used to discern between different CH4 sources and differentiate between the operation of distinct processes [Reeburgh, 2007]. CH4 from biogenic sources tends to have a ∂13CCH4 value of <-50‰ VPDB, while CH4 of thermogenic origin has a ∂13CCH4 value of >-50‰ [Whiticar, 1999]. As light CH4 is preferentially oxidized by microbes, the CH4 pool becomes increasing heavy as it is exposed to oxidative processes. This property can therefore be used to determine whether a decreasing CH4 concentration is due to mixing or dilution, or due to biological oxidation [Reeburgh, 2007]. Previous studies have used this to determine the flow paths of CH4, understand CH4 cycling in ice-surrounded open areas, and identify recently biologically-produced CH4 in the surface ocean [Damm et al., 2005, 2007, 2008; Sasakawa et al., 2008]. Among the various environmental conditions influencing CH4 and N2O cycling, oxygen concentrations are particularly important. As discussed above, low oxygen concentrations favour the production of both gases. Continental shelf seas, particularly those where low oxygen waters are exposed to the surface layers, or those that are underlain by low oxygen sediments, can release large quantities of CH4 and N2O to the atmosphere [Nevison et al., 2004; Naqvi et al., 2010]. Major marine oxygen-deficient zones in upwelling systems, such as those off Peru and Chile and in the Arabian Sea, have been linked to high sea-air fluxes of CH4 and N2O [Bange et al., 1998; Farías et al., 2009]. The Peruvian upwelling system has been shown to contain some of the highest near-surface N2O concentrations in the world, while measured sea-air fluxes of CH4 on the Chilean margin range up to 59 µmol m-2 d-1 [Naqvi et al., 2010 and references therein]. In addition, regimes with abundant supplies of organic matter to the sediments, such as over the Namibian and Indian continental shelves, contain some of the highest near-surface CH4 and N2O concentrations in the open ocean, reaching up to 2900 nmol/L and 436 nmol/L, respectively [Naqvi et al., 2006, 2010; Brüchert et al., 2009]. However, many coastal and low oxygen areas have not been well-studied. High uncertainty in the estimates of global  4 oceanic emissions of CH4 and N2O exists due to the poor sampling resolution in key regions and large ambiguity in sea-air fluxes in many under-sampled areas. 1.2   CH4 AND N2O IN THE ARCTIC OCEAN The Arctic Ocean is thought to be a potentially significant source of CH4 and N2O to the atmosphere, and one that is particularly poorly sampled. The Arctic Ocean’s large, shallow continental shelves are sites of significant CH4 and N2O production and emission. CH4 and N2O super-saturation of up to 125% and ~500%, respectively, has been documented in the surface waters of the Chukchi Sea [Savvichev et al., 2007; Hirota et al., 2009]. The high carbon export in this region is thought to generate CH4 and N2O through the decomposition of organic matter in the sediments [Savvichev et al., 2007; Chang and Devol, 2009; Hirota et al., 2009]. In addition, very high surface concentrations of CH4 in the Eurasian Seas, (mean 18 nmol/L), has been attributed to the thawing of subsea permafrost on these marginal shelves [Shakhova et al., 2010a; Kosmach et al., 2015], while a lesser degree of CH4 super-saturation has been previously found in the permafrost-free Chukchi Sea [Savvichev et al., 2007; Kosmach et al., 2015]. Sea ice also plays a role in regulating these gases by restricting sea-air flux and diluting surface water with low- CH4 and N2O melt water [Randall et al., 2012; He et al., 2013]. Recently, sea ice processes have been identified as potential CH4 and N2O sources associated with microbial activity in anoxic brine channels [Damm et al., 2015a]. In the surface water of the Northwest Passage, the highest CH4 and N2O concentrations were under sea ice, while much lower concentrations were observed in open ocean [Kitidis et al., 2010].  Examination of Arctic CH4 and N2O sources is particularly important in light of the rapid climate-dependent changes occurring in this region. Changes caused by warming alter the biology, chemistry, and circulation of the ocean. Arctic sea ice cover has been drastically thinned and reduced over the past 30 years [Maslanik et al., 2011]. This reduction in sea ice could augment the sea-air flux of CH4 and N2O by lessening exposure to consumption processes [Kitidis et al., 2010], but it may also influence the carbon and nitrogen cycles. The shorter ice cover period and longer growing season has resulted in a 30% increase in primary production in the Arctic Ocean and a northward shift of many species [Arrigo and van Dijken, 2011, 2015; Grebmeier, 2012]. The increased flux of organic  5 matter to the sediments may result in greater CH4 and N2O production due to decomposition processes. The thawing of subsea permafrost may introduce additional organic matter to the system, further increasing the production of these gases, as has been shown in the Eastern Siberian Arctic Seas [Shakhova et al., 2015] and wetland sites in the high Arctic [Elberling et al., 2010]. However, new evidence suggests that efficient anaerobic oxidation of CH4 within the sediments may have the capacity to drastically limit the CH4 flux to the ocean from this source [Overduin et al., 2015]. Warming bottom waters have also caused the destabilization of gas hydrates, which require low temperatures and high pressures to remain stable on the continental slope, and have possibly accelerated the decomposition of the subsea permafrost “cap” which contains hydrates within continental shelf sediment [Ruppel, 2015; Shakhova et al., 2015]. Gas hydrates, although very localized, have been shown to be regionally important sources of CH4 [e.g., Heeschen et al., 2005; Kessler et al., 2006]. The Arctic contains ~100-9000 of Gt CH4 in the form of gas hydrates [James et al., 2016]. The potential destabilization and sudden release of this huge carbon store has attracted much attention due to its association with previous climate change [Dickens et al., 1995, 1997; Lamarque et al., 2006; Archer, 2007; Archer et al., 2009]. Without significant baseline data on the concentrations of CH4 and N2O across varying hydrographic regimes, it is difficult to predict how their concentrations will change into the future. At present, our understanding of the controls on CH4 and N2O cycling in the Arctic Ocean is incomplete. Previous biogeochemical studies in the Arctic Ocean have been spatially segregated [Carmack and Wassmann, 2006], with particular programs focusing on localized regions of interest (e.g. the Bering and Chukchi Seas, Canada Basin, Beaufort Sea, or Davis Strait; Kvenvolden et al., 1993; Hirota et al., 2009; Punshon et al., 2014; Zhan et al., 2015). These studies have shown that CH4 and N2O are frequently supersaturated in the Arctic surface ocean with respect to atmospheric equilibrium values. Studies of Arctic CH4 distributions have focused on the continental shelves, where surface waters are highly supersaturated [e.g., Kvenvolden et al., 1993; Savvichev et al., 2007; Shakhova et al., 2010a], with the greatest attention paid to the high sea-air fluxes in the Siberian Arctic Seas [Shakhova and Semiletov, 2007; Shakhova et al., 2010b, 2013; Kosmach et al., 2015]. By comparison, few studies have sampled the deep basin waters [Damm et al., 2010; Punshon  6 et al., 2014] or in the Canadian Arctic [Kvenvolden et al., 1993; Kitidis et al., 2010; Punshon et al., 2014]. Previous work on N2O has demonstrated high concentrations in the Chukchi Sea [Hirota et al., 2009; Zhang et al., 2015], under-saturation in the deep Canada Basin [Zhan et al., 2015], and low concentrations in sea-ice melt water [Kitidis et al., 2010; Randall et al., 2012; Zhang et al., 2015]. Kitidis et al. [2010] reported measurements of surface water CH4 and N2O concentrations across a broad range of Arctic waters. Their results revealed notable spatial patterns, including 'hot-spots' of high CH4 and N2O under sea ice and in marginal ice zones. These observations are consistent with the recently suggested role of sea-processes as potential CH4 and N2O sources [Damm et al., 2015a], though more depth-resolved observations are needed to further examine this.The work of Kitidis et al. [2010] was largely focused on surface waters (only one depth profile was presented), and thus provided little insight into distribution of CH4 and N2O in sub-surface waters.  The goal of this study was to examine broad-scale, three-dimensional distribution of CH4 and N2O concentrations in the North American Arctic Ocean. During the summer of 2015, depth profiles at nearly 50 locations from the Bering Sea to Baffin Bay were taken in order to quantify the distributions of CH4 and N2O concentrations and sea-air fluxes, and examine potential source and sink terms (Figure 1). To my knowledge, the data set presented here represents the most comprehensive analysis of Arctic CH4 and N2O distributions across a single sampling season. The temporal coherence of the measurements is critical given the rapid changes currently underway in the Arctic. Our results show that N2O is produced in the Bering and Chukchi Seas, and transported with Pacific water masses through to Baffin Bay, with dominant input terms associated with nitrification and denitrification in the Western Arctic. In contrast, CH4 sources to the water column were far more ephemeral, with rapid microbial oxidation of CH4 acting to decrease concentrations and limit eastward transport of CH4. The carbon isotopic composition of CH4 supported this conclusion, as low CH4 concentrations were enriched in 13C due to oxidative processes. This synoptic coverage provides a snapshot of spatial CH4 and N2O variability, and an important baseline dataset for this rapidly changing ocean region.   7 Figure 1: Map of study area, showing the locations of depth profiles (yellow circles) and dominant Pacific Water boundary currents. The black line denotes the transect along which the contour plots (Figure 5 and Figure 10) were derived. Based on [Macdonald et al., 2004].  8 2    Methods and Regional Setting 2.1   FIELD SAMPLING Water samples were collected on three cruises over the period between July to October 2015, covering a transect from the Bering Sea to Baffin Bay (Table 1, Figure 1). At each of the 47 stations, a Sea-Bird CTD system (SBE 911+), equipped with ancillary sensors was used to measure salinity, temperature, and other hydrographic parameters. Discrete water samples for the various stations were collected at multiple depths using 10 L Niskin bottles mounted on a rosette. Salinity values measured by CTD were calibrated against discrete bottle samples analyzed on an Autosal salinometer calibrated to IAPSO standard seawater [Millero et al., 2008]. Table 1: Cruise details Ship Cruise Number Dates Regions Program(s) CCGS Sir Wilfrid Laurier SWL 2015-07 04/07/16 – 26/07/16 Bering Sea, Chukchi Sea C3O, DBO CCGS Amundsen 1502 10/07/16 – 20/07/16 Baffin Bay, Canadian Archipelago ArcticNet, GEOTRACES CCGS Louis St. Laurent LSSL 2015-06 20/09/16 – 17/10/16 Canada Basin, Beaufort Sea JOIS 2.2   DISSOLVED GAS CONCENTRATIONS AND δ13CCH4 VALUES  Water for CH4 and N2O analysis was sub-sampled from Niskin bottles into 60 mL glass vials using thick-walled silicon tubing. Sample vials were overfilled three times and poisoned with 100 µL of saturated HgCl2, taking care to avoid entrainment of bubbles while filling bottles.  Bottles were crimp sealed with no headspace using butyl-rubber stoppers  9 and aluminum caps. Duplicate samples were taken at each depth, and stored at 4°C in the dark until analysis within ~1 month of sampling. CH4 and N2O concentrations were measured simultaneously using an automated purge and trap gas extraction system coupled with gas chromatograph-mass spectrometer (PT-GCMS) at the University of British Columbia according to the method of Capelle et al. [2015]. The pooled standard deviations of the duplicates were 0.88 and 0.58 nmol/L for CH4 and N2O measurements, respectively. As discussed by Capelle et al. [2015], the accuracy and precision of our automated system compares very well against methods by other leading laboratories around the world.  Percent saturation relative to atmospheric equilibrium was calculated from the mean concentration of the duplicate samples and gas solubility was calculated following Wiesenburg and Guinasso Jr. [1979] and Weiss and Price [1980] for CH4 and N2O, respectively, using the CTD measured in situ temperature and salinity. Atmospheric concentrations of CH4 and N2O were taken as 1910 ppb and 328 ppb, respectively, based on the average mixing ratios in Barrow, Alaska in August 2015 (Data provided by NOAA ESRL Global Monitoring Division, Boulder, Colorado, USA; http://esrl.noaa.gov/gmd/). Samples for δ13CCH4 analysis were collected in either 160 mL or 250 mL glass vials in the same manner as above, and poisoned with either 300 µL or 500 µL HgCl2, respectively. The values were determined using a Delta XP plus Finnigan mass spectrometer at the Alfred Wegner Institute in Bremerhaven, Germany. The samples were pre-concentrated through purging and subsequent trapping with PreCon equipment (Finnigan), and the isotopic ratios were given in delta notation relative to the Vienna Pee Dee Belemnite (VPDB) standard. 2.3   NUTRIENT ANALYSIS For cruise LSL 2015-06, unfiltered nutrient samples from were analyzed for nitrate + nitrite, phosphate, and silicate on a three channel Seal Analytical Nutrient Auto-Analyser 3 (AA3) within 12 hours of collection. For cruise SWL 2015-07, nutrient samples were analysed by the same method, but were frozen prior to being analyzed at the Institute of Ocean Sciences 5 months after the cruise. Unfrozen ammonium (NH4+) samples were analyzed by the Nutrient Analytical Services Laboratory at the University of Maryland, following the method described in Parsons et al. [1984].   10 2.4   BIOGEOCHEMICAL CALCULATIONS N* is a measure of the relative excess or deficit of fixed nitrogen relative to phosphate, calculated using the equation N* (µmol/L) = ([NO3-] + [NO2-] +[NH4+] -16*[PO43-] + 2.9 of [Gruber and Sarmiento, 1997]. Assuming Redfield stoichiometry (C:N:P = 106:16:1), a negative N* value indicates a deficit of dissolved inorganic nitrogen, while a positive value indicates N fixation [Redfield, 1942; Gruber and Sarmiento, 1997]. Apparent oxygen utilization (AOU) represents the amount of oxygen that has been consumed by remineralization since the water mass was last ventilated. AOU is calculated as difference between the dissolved oxygen concentration measured in a sample and the calculated concentration at equilibrium with the atmosphere [Redfield, 1942].  For the purposes of these calculations, we (and others) assume that temperature and salinity characteristics of waters masses have not been altered since they were last in contact with the atmosphere.   2.5   SEA-AIR FLUX CALCULATIONS Sea-air fluxes of CH4 and N2O (µmol m-2 d-1) at each station were calculated using a mixed layer depth computed based on a density difference criterion of 0.125 kg m-3 relative to 3m depth. If no measurement was taken within this calculated mixed layer, measurements within 3 m of the mixed layer depth were used instead for flux calculations. The disequilibrium of CH4 and N2O in the mixed layer was calculated by subtracting the atmospheric saturation value (by the same method described in section 2.2) from the observed concentration. Wind data were acquired from the NCEP Reanalysis data provided by the NOAA/OAR/ESRL PSD, Boulder, Colorado, USA, from their website at http://www.esrl.noaa.gov/psd/. The equation for the ocean-atmosphere flux was taken as the product of a gas exchange coefficient, kw, and mixed layer disequilibrium, ΔC: Flux = kw × ΔC The gas exchange coefficient (kw) was calculated as a function of the Schmidt number (Sc) and wind speed according to the parameterization by Nightingale et al. [2000]. This equation was developed for ice-free conditions, which is appropriate for our calculations since we did not have any mixed layer gas samples for ice-covered stations. A weighting function [Reuer et al., 2007] was used to incorporate the effects of variability in wind speeds  11 in the 60 days prior to sample collection.  The net result is a time-weighted piston velocity, where the weighting is based on the fraction of the mixed layer ventilated on any given day.   2.6   REGIONAL SETTING In order to provide a context for our CH4 and N2O concentration data, we briefly review the dominant oceanographic properties of our broad survey region.  We sampled across a number of distinct water masses, which are influenced by a range of physical and biological processes that impact salinity, temperature, and nutrient concentrations.  These tracers allow us to better constrain the biogeochemical processes that influence trace gas cycling.  We also briefly review the available information on CH4 emissions from seafloor sediments across our survey region. 2.6.1   Bering and Chukchi Seas The Bering and Chukchi Seas are dominated by northward flowing Pacific water that comprises several distinct water masses. As these water masses transit over the shallow Bering and Chukchi Shelves into the Arctic Basin, they are modified by biological and physical processes [e.g., Yamamoto-Kawai et al., 2006]. In the northern region of the Bering shelf, three main water masses are present in the summer. The Anadyr Water occupies the western portion of the Bering Sea, and is composed of upwelled, saline, nutrient-rich deep water in the Gulf of Anadyr. The eastern Bering Sea is dominated by relatively fresh (S <31.8), nutrient-poor (<1.5 µmol/L NO3-) Alaskan Coastal Water (ACW). A central water mass with intermediate salinity is known as Bering Sea Water (BSW). Mixing of the BSW and Anadyr Water is enhanced in the Northern Bering Sea and Bering Strait, but some unaltered Anadyr Water and ACW persists into the Chukchi Sea [Weingartner et al., 2005; Grebmeier et al., 2006]. The BSW contains high dissolved nutrient concentrations due to the influence of the Anadyr Water, but seasonal differences in sea ice growth results in two distinct water masses, known as Pacific Winter Water (PWW) and the Pacific Summer Water (PSW). In the winter, convective overturning caused by brine rejection stirs up the sediments and adds regenerated sedimentary nutrients to the PWW [Pickart et al., 2016], while imparting a low temperature and high  12 salinity signature [e.g., Woodgate et al., 2005a; 2005b]. During summer, the fresher, northward flowing PSW gradually displaces the PWW on the Chukchi shelf, although remnant PWW is still present on the northern Chukchi Shelf [e.g., Lowry et al., 2015; Pickart et al., 2016]. At the northern-most point in our study area, in Barrow Canyon, all of these water masses converge and stratify [Gong and Pickart, 2015]. The dense PWW sinks to the deepest parts of the canyon (below 35m depth) and is overlain by PSW and ACW (5 - 35 m) as well as a surface melt-water layer (0 - 5 m) [Gong and Pickart, 2015].   The contrasting nutrient concentrations in the different water masses affect primary production, which has important implications for CH4 and N2O cycling. The high nutrient BSW sustains an annual productivity of up to ~ 500-700 g C m-2 y-1 in the northern Bering and southern Chukchi Seas, with values declining to ~ 50-100 g C m-2 y-1 in the northern Chukchi Sea [Hameedi, 1978; Springer and McRoy, 1993]. By contrast, the nutrient-poor southeastern Chukchi Sea supports primary production of less than 100 C m-2 y-1 [Springer and McRoy, 1993]. This east-west productivity gradient is also reflected in organic carbon fluxes and benthic oxygen consumption rates, which are both significantly higher in regions of elevated primary productivity [Naidu et al., 2004]. The high oxygen consumption rates in the productive regions result in sediment anoxia, creating conditions conducive to anaerobic microbial processes associated with CH4 and N2O production. Indeed, denitrification has been well documented on the Bering and Chukchi shelves [Devol et al., 1997; Tanaka et al., 2004; Yamamoto-Kawai et al., 2006], and methanogenesis has also been reported [Savvichev et al., 2007]. These processes could add CH4 and N2O to the water column in the Bering and Chukchi, which could potentially be transported to other regions of the Arctic by eastward flowing waters.  Rivers are another potential source of CH4 and N2O to surface waters of the Western Arctic, delivering CH4 released from terrestrial permafrost, and CH4 and N2O produced in riverine sediments to coastal regions.  The Yukon River, whose drainage basin contains near-surface permafrost [Pastick et al., 2014], is the second largest river in the North American Arctic [Cooper et al., 2008]. Yukon river water feeds directly into the ACW, and high measured concentrations of CH4 (up to 330 nmol/L) in this river and its tributaries have been attributed to the anoxic decomposition of the organic matter in its watershed [Striegl et al., 2012]. We are unaware of any N2O data for the Yukon River.    13 2.6.2   Canada Basin and Beaufort Sea The Canada Basin comprises deep waters (~4000 m) bordered by shallow seas and continental shelf. Surface waters of the Canada Basin are characterized by a ~ 40 m thick Polar Mixed Layer (PML) that is relatively fresh (26.2-31.2 ‰) and nutrient-depleted [Macdonald et al., 2004]. As Pacific waters enter the Canada Basin from the Chukchi Sea, they subduct below the PML into the Upper Halocline Layer (UHL) [Codispoti et al., 2009].  Within the UHL, Pacific Summer Water (PSW) (S= 31–32, T >-1°C, D= 25-26 σt) is found between ~40-100m, while Pacific Winter Water (PWW) (S= 32–33.1, T <-1°C, D= 26-26.5 σt) sits between ~100-200m. Previous authors have used a salinity of 33.1 and temperature of -1.6°C as the signature characteristics of the PWW core [e.g., Aagaard et al., 1981; Jones and Anderson, 1986; Carmack, 2000]. The ACW is only a very minor component of the UHL, and is generally considered part of the PSW once it enters the basin [Shimada et al., 2001]. Below the UHL, the Lower Halocline Layer occupies a depth range of ~200-400 m and is largely of Atlantic origin, while Canada Basin Intermediate Water (CBIW) occupies a depth range of ~ 400-2000m. Canada Basin Deep Water (CBDW), which is found below ~2000m, has been isolated from the atmosphere for ≈ 450 years [Schlosser et al., 1997; McLaughlin et al., 2002; Timmermans et al., 2003]. Strong salinity stratification in the surface waters of the Canada Basin results in low surface water nutrient concentrations and, thus, restricted primary productivity [Sakshaug, 2004].  This, in turn, results in low carbon export to sub-surface waters, which have been isolated from the atmosphere for decades or longer [Macdonald et al., 2004]. As a result, the chemistry of deep basin waters is relatively homogenous [Timmermans et al., 2003]. The Beaufort Sea, in the southern Canada Basin, is primarily influenced by the Upper Halocline water masses of the Canada Basin, including both PSW and PWW [Carmack et al., 2006].  Additionally, the ACW wraps around Barrow Point and is transported into the Beaufort Sea. This inflowing ACW is retained close to the coast, and was not likely sampled during our surveys. The carbon export in the ACW inflow is considerably lower than that of the Bering and Chukchi Seas (due to lower nutrient concentrations), resulting in lower benthic productivity and oxygen consumption [Naidu et al., 2004].    14 2.6.3   Canadian Arctic Archipelago The Canadian Arctic Archipelago (CAA) is a vast stretch of islands and channels that lies to the north-east of the Beaufort Sea (Figure 1).  This region receives Pacific-derived water from the west, with PML and UHL waters flowing eastward over the sill demarcating the entrance to the Amundsen Gulf. Westward flow is limited by a shallow sill near Resolute, such that almost all of the water in the CAA (even as far east as Lancaster Sound) is derived from Pacific-origin water masses with only a limited influence of Atlantic-derived waters [Jones, 2003; Michel et al., 2006]. The complex network of shallow, narrow straits in the CAA is subject to significant regional and small-scale variability. River input modifies the inflowing water, and imparts a low salinity signature on the southern CAA [McLaughlin et al., 2004]. Local mixing and tidal processes through the narrow straits and channels also reduce stratification and increase nutrient concentrations [McLaughlin et al., 2004, 2011]. High sea ice cover also influences biogeochemical cycling in the CAA. Sea ice acts as a semi-permeable barrier to sea-air exchange of dissolved gasses [e.g., Loose et al., 2011], and has been suggested as an explanation for the accumulation of high CH4 and N2O concentrations in the surface waters of the CAA [Kitidis et al., 2010; Damm et al., 2015a]. Primary productivity estimates in the CAA vary from 10 to 55 g C m-2 y-1 [Welch et al., 1992; Sakshaug, 2004]. Sea ice algae are an important part of this production, contributing 5 to 10 g C m-2 y-1 [Smith et al., 1988; Michel et al., 2006]. The linkages between primary productivity and benthic remineralization have not been well-studied within the CAA, but recent data have shown relatively high benthic remineralization rates in Lancaster Sound, and much lower rates in the Amundsen Gulf and Baffin Bay [Darnis et al., 2012]. Regions with high remineralization rates may be sources of CH4 and N2O to the CAA. 2.6.4   Seafloor methane sources The shallow shelves of the Arctic Ocean were once vast coastal plains which developed layers of thick permafrost during the Late Pleistocene (Brigham and Miller, 1983). Gas from petroleum deposits or biological production deep in the crust migrated upwards and was frozen in place as gas hydrates [Ruppel, 2015]. A post-glacial sea level rise of 120m since ~12ka inundated these plains, thawing the permafrost deposits and likely causing the  15 dissociation of permafrost-associated gas hydrates [Ruppel, 2015]. This release of gas in unstable hydrates, either entrained in or capped by a permafrost layer, are thought to be the sources of the large methane emissions on the wide Eastern Siberian Arctic Shelf and other Eurasian shelves [Shakhova et al., 2010b, 2015]. In contrast, the North American Arctic only holds permafrost deposits (extending to ~100m depth) and shallow gas hydrates on the narrow Beaufort Shelf [Paull et al., 2007, 2011; Ruppel, 2015]. Widespread gas plumes have been documented on the continental margin of the Beaufort Sea, and these have been attributed to decomposing permafrost and gas hydrates [Paull et al., 2011]. Deepwater gas hydrates (~300-1200m) are also present on the Beaufort Slope [Weaver, J.S., Stewart, 1982; Dallimore, 1999; Ruppel, 2015]. These hydrates on the continental slope are generally considered the most sensitive to disruption due to climate change [Kvenvolden, 1993; Ruppel, 2011]. Warming over the past 60 years has been implicated in the potential destabilization of gas hydrate deposits between 300 and 550 m depth [Phrampus et al., 2014]. Additionally, a survey of the slope sediments has shown elevated sediment-ocean fluxes of CH4 from 280-1500 m depth on the western Beaufort slope, potentially attributable to gas hydrate dissociation [Coffin et al., 2013]. Hydrocarbon deposits may also play a role as a CH4 source. Much of the Chukchi Sea and Canada Basin, including the Beaufort Sea, is underlain by hydrocarbon-bearing source rocks [Burlin and Shipel’kevich, 2006; Grantz and Hart, 2012]. These can contribute CH4 to the water column though upward migrations along faults and fractures in the rock [Reeburgh, 2007]. Thermogenic gas seepages are widespread in areas where oil and gas deposits occur, and are often used as targets of commercial exploration [Hovland et al., 1993]. However, there is only limited published information on the exact location of such seepage sites [Hovland et al., 1993]. As the Chukchi and Beaufort Seas have been extensively explored for offshore oil and gas, it is likely that thermogenic seepages exist in these areas and contribute to the CH4 cycling.  16 3    Results and Discussion For the purpose of describing gas distributions, we divide our study area into three sub-regions: a) Bering Sea and Chukchi Sea, b) Canada Basin and Beaufort Sea, and c) Canadian Arctic Archipelago and Baffin Bay. This grouping was based on the similarity of hydrography and CH4 and N2O distributions within these regions. We first provide an overview of CH4 and N2O concentrations across our full transect, and then focus on the processes occurring in each of these sub-regions.  Since the cycling of CH4 and N2O is influenced by distinct biogeochemical processes, we discuss each of these gases separately. 3.1    NITROUS OXIDE  Nitrous oxide concentrations across the entire transect ranged from 10.9 to 24.6 nmol/L, with corresponding saturation levels of 77% to 145%. As shown in Fig. 2, concentrations of N2O exhibited a ubiquitous maximum in the Pacific Winter Water (PWW), with persistent low values in the surface fresh-water layer and ACW (Figure 2). This feature can be seen in Fig. 3, where the average profiles of N2O from the Bering, Chukchi and Beaufort Seas, and the CAA region are plotted on density surfaces. As discussed below, this high N2O signal in the PWW originates in the Bering and Chukchi Seas and is transported across the Arctic Ocean on the 26.5 σt density surface until it is attenuated in the Canadian Arctic Archipelago. In deeper waters, N2O concentrations tend to decrease with increasing density, while N2O concentrations in the lower-density surface waters are more variable because they are influenced by sea-ice cover, water mass mixing, and sea-air exchange.  17 -2024681019 20 21 22 23 2424252526262728N2O (% saturation)85909510010511011512012519 20 21 22 23 242425252626272824 26 28 30 32 34Temperature (°C)-20246810CH4 (% saturation)25 40 63 100160250400630Salinity (‰)PSWPWWACWMelt-water / River inputAWPSWPWWACWMelt-water / River inputAWFigure 2: Salinity vs. temperature plot showing the different water masses observed in the North American Arctic. The top panel shows percent saturation of N2O, while the bottom panel shows percent saturation of CH4. Circles, squares, and triangles denote measurements from the Bering and Chukchi Seas, Canada Basin, and Canadian Arctic Archipelago, respectively.  18  Figure 3: Density vs. concentration plots averaged by sub-region. The grey shaded area is the standard deviation of the mean. 3.1.1   N2O in the Bering and Chukchi Seas Across all sampling depths, N2O concentrations ranged from values just under atmospheric equilibrium in the Northern Bering Sea (11.8 nmol/L) to a maximum of 24.6 nmol/L (145% saturation) at the northernmost station in the Chukchi Sea (Figure 4).  Surface water N2O concentrations were low, and typically less than the atmospheric equilibrium concentration of ~13 nmol/L. These concentrations are consistent with observations reported previously for this region (~11–25 nmol/L) [Hirota et al., 2009; Zhang et al., 2015]. The nearly ubiquitous under-saturation of N2O in surface waters may reflect dilution by sea ice meltwater which contains little N2O [Randall et al., 2012].  In contrast to previous observations in our study area [Hirota et al., 2009; Zhang et al., 2015] we did not detect high (>120%) surface super-saturations of N2O in the Bering and 12 14 16 18 202022242628BERING SEAN 2O (nmol L  )12 14 16 18 20CHUKCHI SEA12 14 16 18 20CANADA BASIN12 14 16 18 20CANADIANARCHIPELAGO5Density (σt)2022242628CH4  (nmol L  )-1-110 15 20 5 10 15 20 5 10 15 20 5 10 15 20 19 Chukchi Seas. The lower concentrations we observed may reflect a difference in the geographic region and/or seasonal timing of our sampling compared to previous studies. Hirota et al. [2009] sampled stations further to west, with significant influence of high-nutrient Anadyr Water, which likely fuels high productivity and oxygen depletion in the sediments leading to increased N2O concentrations.   Below the surface layer, N2O distributions appeared to be influenced by local water masses and biological processes in the sediments. In particular, differences in nutrient availability and biological productivity across distinct water masses appeared to influence N2O accumulation as water transited through the Bering and into the Chukchi Sea. The northern Bering Sea contained low and near atmospheric equilibrium N2O concentrations throughout the water column, averaging 14.4 nmol/L (Figure 4). In the southern Chukchi Sea, two separate patterns emerge, with differences in N2O concentrations in the ACW and PSW likely related to contrasting productivity and carbon export in these waters [e.g., Naidu et al., 2004].  Some of the lowest N2O concentrations along the transect were measured in ACW waters  that were characterized by low nutrient concentrations and biological productivity (stations UTN-1, SEC-5, SEC-7, Figure 4). This low productivity results in lower carbon fluxes [Naidu et al., 2004], limiting denitrification and nitrification sources of N2O. In contrast, the PSW-influenced central channel waters were characterized by higher nutrient concentrations and phytoplankton biomass, and high benthic oxygen consumption has been previously measured in this region [Grebmeier, 1993]. As expected, we observed higher N2O concentrations in the central channel (stations UTN-3, UTN-7, SEC-1, SEC-3).  20  Figure 4: Map and corresponding N2O depth profiles in the Bering and Chukchi Seas. Black filled circles denote a measurement, while the error bars represent the standard deviation of the duplicate samples. The grey dotted line is the atmospheric equilibrium concentration.  The highest N2O concentrations along the entire transect (24.6 nmol/L N2O) were measured at the northernmost stations (BarC-8) in the remnant PWW at a depth of 65 m. On average, the N2O concentration below the shallow (5-15m) fresher water surface layer was 18.0 nmol/L, representing an increase of 3.6 nmol/L relative to the PSW. Similarly, the accumulation of N2O in the bottom waters of the northward-flowing Bering and Chukchi Seas was also observed by Hirota et al. (2009). To further illustrate this latitudinal gradient in N2O concentration, Fig. 5 shows a latitude-depth contour plot highlighting the strong north-south N2O gradient along the 26.5 isopycnal (S≈33) (Figure 5). This pattern may reflect the accumulation of N2O in the bottom waters as they flow over the Chukchi Shelf.  21 An additional explanation for the high N2O concentrations is the slower circulation rates of PWW over the shelf [Woodgate et al., 2005a], which would result in an increased time for N2O accumulation in the winter water, as denitrification rates remain constant throughout the year [Devol et al., 1997; Chang and Devol, 2009]. Furthermore, brine rejection during ice growth in the winter generates convective overturning, stirring the sediments and likely adding N2O to the PWW [Pickart et al., 2016].   Figure 5: Nitrous oxide and density contour plots. The top two panels show the distribution of N2O and density along a ~1500km transect from the Bering to the Beaufort Sea (black line on Fig. 1), respectively. The bottom panel shows the mean concentration of N2O on the 26.5 density surface, illustrating the accumulation of N2O along this isopycnal. To determine the sources of N2O in the Bering and Chukchi Sea regions, we examined the relationship between N2O concentrations and other hydrographic variables. Water column N2O production in these shallow regions may be limited, as nitrifying bacteria are potentially light-inhibited [Ward, 2008], and water column oxygen concentrations are too high for denitrification. Rather, sedimentary processes likely represent the predominant N2O source to the water column, and this is consistent with the observation of maximum along the 26.5 σ  surfacetDistance (km)0 500 1000 1500N2O (nmol L-1)14161820-120-100-80-60-40-200Depth (m)-120-100-80-60-40-2002424.52525.52626.523.523Density (t)62 64 66 68 70Latitude (°N)151617181920141312σ11N2O (nmol L-1 )Bering Sea Chukchi SeaBering StraitBarrow Canyon 22 concentrations in the bottom waters (Figure 4). Previous studies have attributed the high N2O concentrations on the Bering and Chukchi Sea shelves to denitrification in the sediments [Hirota et al., 2009]. Indeed, we observed a negative correlation between ΔN2O and N* (Figure 6) in this region, providing evidence that denitrification is a source of N2O to the water column.  Although the relationship between N* and N2O is statistically significant (p<0.0001), it exhibits significant scatter (r2 = 0.18).  Some of this variability is likely attributable to the influence of other nitrogen cycling processes in N2O production.  Recently, nitrification-denitrification coupling has been identified as an important process on the Bering and Chukchi Sea shelves [Granger et al., 2011; Horak et al., 2013; Brown et al., 2015]. This pairing has been shown to be an important mechanism in N2O production [Naqvi et al., 1998]. Typically, the correlation of ΔN2O with AOU is used to determine if nitrification is a source of N2O. In our data set, we found no significant relationship between the two variables.  This result may be attributable to large differences between the oxygen concentration in the water column (>230 µmol/L O2) compared to the sediments (due to high benthic O2 consumption; Grebmeier et al., [2006b]), such that the water column AOU does not reflect the conditions of N2O production.  Although we cannot determine the exact mechanism of N2O formation, it is apparent that the active nitrogen cycling on these shelves is the predominant source.  23  Figure 6: ∆N2O vs. AOU and N* plots, broken down by sub-region. Only the Canada Basin UHL is shown in order to better compare the properties of the Pacific water masses between regions. Trend lines are shown where there is a statistically significant regression model (p<0.05).  3.1.2   N2O in the Canada Basin and Beaufort Sea Nearly all the N2O profiles in the Canada Basin and Beaufort Sea follow the same pattern, with under-saturated surface water concentrations (91-98% saturation), increasing to a maximum in PWW, then decreasing to a minimum in the CBDW (~75% saturation) (Figure 7). The high N2O signature of the Bering and Chukchi seas is transported with the Pacific Winter Water into the Canada Basin and Beaufort Sea. The pronounced N2O maximum is tightly constrained to values ranging from 18.5-19.6 nmol/L, and it occurs at S~ 33‰ (26.5 σt), in the core of the PWW (Figure 3).  Bering and Chukchi Seas-2-101234 R2 = 0.2%p = 0.58R2 = 18.2%p<0.0001∆ N2O (nmol L-1)-2-101234 R2 = 46.2%p<0.0001R2 = 46.1%p<0.0001AOU (µmol L-1 )-50 0 50 100 150-2-101234 R2 = 0.71%p = 0.93N* (µmol L -1)-15 -10 -5 0R2 = 13.6%p<0.0001Canada Basin and Beaufort SeaCanadian Arctic Archipelago 24  Figure 7: Map and N2O depth profiles in the Canada Basin and Beaufort Sea. We observed a positive correlation of ΔN2O and AOU in the Canada Basin samples (Figure 6).  In previous studies, this relationship has been taken as evidence of a primary nitrification source for N2O in seawater [Nevison et al., 1995].  However, several lines of evidence suggest that the nitrification is only a minor source of N2O in the Canada Basin.  Based on the slope of the AOU vs. ΔN2O in the UHL of the Canada Basin (40-200m), we calculate 0.016 nmol of N2O produced for every µmol of O2 consumed.  Assuming Redfield stoichiometry, with 17% of oxygen consumption attributed to NH4+ oxidation [Ward, 2008], we derive an N2O yield from nitrification of 0.01%. This is value falls within the range previously observed for waters of the Subarctic Pacific Ocean (0.004-0.028%; [Yoshida et al., 1989; Grundle et al., 2012]). Previous studies have reported a ~1.4 µmol/L difference between the NH4+  concentration on the Chukchi shelf and that in the Canada Basin [Zhang et al., 2015]. The nitrification of 1.4 µmol/L NH4+ would therefore produce 0.14 nmol/L N2O, which represents approximately 7% of the observed N2O surplus in the Canada Basin PWW (~2.0 nmol/L). Even assuming nitrification of up to 6 µmol/L NH4+ (the highest concentration of NH4+ documented on the Chukchi shelf [Brown et al., 2015]), only  ~ 0.6 nmol/L N2O would be produced. This suggests that the high N2O signature in the Canada Basin is set by processes occurring in the Bering and Chukchi shelves, with  25 only a negligible amount added by in situ water column nitrification within the Canada Basin. This notion is supported by the negative correlation of ΔN2O with N* in the UHL of the Canada Basin, which implies that denitrification is a major source of N2O (Figure 6), similar to the relationship we observed on the Bering and Chukchi Sea shelves. Previous authors have raised concerns about the use of N2O-AOU relationships as an indicator of nitrification as a primary N2O source [Nevison et al., 2003; Bange, 2008]. In particular, Yamagishi et al. [2005] showed that denitrification-sourced N2O assumed a linear relationship with AOU when diffusing into deep waters,. This implies that a linear relationship between N2O and AOU can result from water mass mixing rather than remineralization processes. The Beaufort Sea is strongly influenced by the upper layers of the Canada Basin. Our observations point to physical processes, rather than microbial ones, as key determinants of N2O distributions.  As an example of this, our results demonstrate the regional importance of upwelling on N2O distributions. This was especially evident at CB28aa, where PWW from MK1 upwelled onto the shelf, replacing the water at this shallow shelf station with PWW and yielding high N2O concentrations in shallow (<60m) waters (Figure 7). Deep Canada Basin waters contain low and relatively constant concentrations of N2O, with an average value of 13 nmol/L (80% saturation).  Our observations are in good agreement with those of Zhan et al. [2015], who observed an average concentration of 12.5 nmol/L in these waters. These authors attributed the large apparent under-saturation to the historical characteristics of the deep water. Deep Canada Basin waters were last exposed to the atmosphere ~450 years ago [Schlosser et al., 1997], when atmospheric N2O concentrations were 16% lower than modern values [Flückiger et al., 1999]. Using the equation by Weiss and Price [1980], and assuming an atmospheric mixing ratio of 270 ppb for the year 1700 [Flückiger et al., 1999], the equilibrium concentration at that time was ~13 nmol/L, which is equal to our average concentration in the deep Canada Basin waters. This provides an explanation for the strong apparent under-saturation of these waters when expressed relative to modern equilibrium values.  26 3.1.3   N2O in the Canadian Arctic Archipelago Surface concentrations of N2O in the CAA varied from 13.1 nmol/L (95% saturation) to 17.8 nmol/L (110% saturation). Concentrations greater than atmospheric equilibrium were generally observed in association with ice cover (CAA-4, CAA-7, VS), and a maximum of 16.0 – 18.5 nmol/L was observed in the Pacific Winter Water (26.5 σt) (Figure 8). These maximum concentrations were lower than those measured in the Canada Basin. Below the N2O maximum, concentrations generally decreased with depth to near atmospheric equilibrium values (~14.5 nmol/L). We observed a notable decrease in the maximum N2O concentration east of the Amundsen Gulf, with the exception of stations AN-312 and AN-314, which were too shallow to contain PWW. Similar to the other regions, we observed a negative correlation between ΔN2O and N* (Figure 6), implying that the N2O maximum in the CAA is originates from the Pacific shelves. This result suggests that the N2O signal is being transported thousands of kilometers across the Arctic Ocean along the 26.5σt  isopycnal, until it is attenuated in the CAA (Figure 3). The reduction in N2O concentrations in the CAA can be attributed to sea-air flux (see below), and the mixing of the PW with low N2O sea ice melt-water and/or river water. In Baffin Bay, N2O concentrations were close to atmospheric equilibrium values through much of the water column, increasing with depth to a sub-surface maximum at 1500 m of 18.0 nmol/L (115% saturation) (Figure 8). A similar deep N2O maximum in Baffin Bay was reported by Kitidis et al. [2010], who attributed this feature to sedimentary denitrification in the sediments just off of Greenland [Rysgaard et al., 2004].  27  Figure 8: Map and depth profiles of N2O in the CAA. 3.2   METHANE Across our entire sampling transect, CH4 concentrations varied greatly, ranging from 0.7 to 30.5 nmol/L, corresponding to percent saturation of 30% to 800%. We observed the highest concentrations of CH4 in the remnant PWW of the northern Chukchi Sea (Figure 9). Unlike N2O, CH4 was oxidized relatively rapidly in the water column, limiting the transport of high CH4 concentrations in eastward flowing waters (Figure 3). As a result, the high CH4 concentrations present in the Chukchi Sea, at a density surface of 26.5 σt, were not observed in either the Beaufort Sea or CAA region.  In the Canada Basin, a ubiquitous and stable maximum in the UHL associated with the Pacific Summer Water (PSW) was measured at 25.5 σt (Figure 3). By contrast, CH4 concentrations below the UHL varied  28 greatly, ranging from 1 to 24 nmol/L. CH4 concentrations in the Canadian Arctic Archipelago (CAA) showed large surface maxima, up to 15.0 nmol/L, and decreasing concentrations with depth to near atmospheric equilibrium values. This spatial and temporal variation underscores the ephemeral nature of CH4 sources and sinks. 3.2.1   CH4 in the Bering and Chukchi Seas Methane concentrations in the Bering and Chukchi Seas were almost ubiquitously supersaturated throughout the water column. Concentrations ranged from slightly greater than atmospheric equilibrium in the surface of the Bering Sea (~3 nmol/L) to values close to 800% saturation (30.5 nmol/L) in the northern Chukchi Sea (Figure 9). CH4 distributions appeared to be more influenced by local sources than N2O, including potential sources from rivers feeding into the ACW, and sediment-derived CH4 in the northern Chukchi Sea water column. Surface water CH4 concentrations were frequently supersaturated, averaging 4.8 nmol/L (150% saturation). No spatial concentration gradient was discernable within our study area. The surface water CH4 concentrations we measured were lower than the average surface CH4 concentrations previously reported for the Chukchi Sea by Kosmach et al. [2015] and Savvichev et al. [2007] (8.1 nmol/L and 11.2 nmol/L, respectively). These previous studies sampling further west in the Chukchi Sea, where higher CH4 concentrations have previously been reported relative to the eastern sections [Savvichev et al., 2007].  29  Figure 9: Map and depth profiles of CH4 in the Bering and Chukchi Seas Below the surface mixed layer CH4 concentrations increased considerably from the northern Bering Sea (mean ~5 nmol/L) to the northern Chukchi Sea (mean ~17 nmol/L) (Figure 10). Our measurements from just south of the Bering Strait contained increasing concentrations of CH4 with depth (up to ~7.0 nmol/L, 210% saturation). In the southern Chukchi Sea, elevated CH4 concentrations (mean 8.7 nM, range: 5.3-12.3 nmol/L) were observed in the ACW (stations UTN-1, SEC-5, SEC-7) while lower CH4 concentrations (mean 6.3 nmol/L, range: 3.7-10.1 nmol/L) were measured in the more PSW-influenced western stations (UTN-3, UTN-7, SEC-1, SEC-3).  This is somewhat surprising given that rates of methanogenesis in the sediments of the southern central Chukchi Sea have been found to be five times greater than those near the Alaskan Coast [Savvichev et al., 2007]. Instead, the greater concentrations of CH4 in the ACW may reflect the high CH4  30 concentrations associated with river inputs into the ACW, which results from the methanogenesis of organic material in degrading permafrost [Striegl et al., 2012]. This implies that methanogenesis in oceanic sediments is perhaps only minor compared to the contribution of CH4 by rivers.  Figure 10: Methane and density contour plots. The top panel shows the distribution of CH4 along a ~1500 km transect from the Bering to the Beaufort Sea (black line on Fig. 1). The bottom panel shows the density distribution along the same transect. The highest dissolved CH4 concentrations in the study area were observed in the sea ice-influenced remnant PWW in the Northern Chukchi Sea (30.6 nmol/L) near the sediments. Previous work has demonstrated high concentrations of CH4, averaging ~2200 nmol/L, in the upper 2 cm of sediments of the northern Chukchi Sea (Lapham et al., submitted). Brine rejection, as a result of sea ice growth, causes convective overturning of the water column and mixes the sediments, releasing regenerated nutrients and gases into the PWW [Pickart et al., 2016]. As this process only occurs in the winter, it may account for the higher CH4 concentrations in the PWW compared to the PSW. The origin of seafloor CH4 sources is unclear. As this area does not contain subsea permafrost or gas hydrates [Ruppel, 2015], it is uncertain whether the elevated CH4 signature results from methanogenesis in the sediments or perhaps upward migration of along the 26.5 σ  surfacettDistance (km)0 500 1000Bering Sea Chukchi SeaBering StraitBarrow Canyon-120-100-80-60-40-200Depth (m)-120-100-80-60-40-201012141618208642424.52525.52626.523.523Density ( σt)CH4  (nmol L  )-162 64 66 68 70Latitude (°N)01500 31 thermogenic gas. Rates of methanogenesis of up to 67 µmol m-2 d-1 have been measured in the sediments of the Chukchi Sea [Savvichev et al., 2007], implying that this may be a major source of the water column CH4. This area is also in the Chukchi Sea Oil and Gas Lease Sale No. 193, suggesting likely geologic gas deposits.   In addition to the elevated CH4 concentrations in bottom waters of the Chukchi Sea, we also observed CH4 maxima within the upper water column at a few stations (Figure 9) (DBO4.3, DBO4.2, BarC-10, BarC-6, BarC-4).  These signatures may be due to in situ methane production within the water column. The most likely processes are methanogenesis in anoxic microenvironments [de Angelis and Lee, 1994; Holmes et al., 2000] or the cleavage of methane from DMSP [Damm et al., 2010]. The cleavage of methyl groups from methanophosphates is unlikely in this case because this region is not phosphate limited [Karl et al., 2008]. Methane has been found to be produced in the ocean where suspended sediments accumulate, presumably in the anaerobic niches provided by these particles [Holmes et al., 2000]. However, we did not find any relationship between the transmissivity of the water column (a proxy for suspended particle loads) and CH4 concentrations. The second possible pathway is the bacterial metabolism of DMSP, a compound produced by phytoplankton, producing precursors for methane formation. A methyl group may be cleaved when bacterioplankton use DMSP as a source of carbon, which occurs predominantly when they are nitrate-limited [Kiene et al., 2000; Damm et al., 2010]. Indeed, we did observe see some co-occurrence of high CH4 with high chlorophyll in this region (Figure 13), implying a potential phytoplankton-derived source of DMSP, could stimulate CH4 release during bacterial DMSP metabolism. In the absence of sub-surface DMSP data for this region, we cannot conclusively link DMSP metabolism to CH4 production.   Stable isotopic composition of CH4 in the Bering and Chukchi Seas The large range of our CH4 concentrations in our data suggests that sources of CH4 are variable, and processes that both consume and produce CH4 are pervasive. We used the stable isotope composition of CH4 to provide insight into CH4 cycling in the water column. Some of the lightest ∂13CCH4 values (~-41.5‰) were found in the bottom waters of the northern Chukchi Sea, associated with high concentrations of CH4 (up to 31 nmol/L). This  32 -41.5‰ minimum is indicative of a more recent source. Although it is relatively heavy for methane produced by microbial methanogenesis (normally considered <-50‰ [Whiticar, 1999]), this may reflect the oxidation of biogenic methane within the sediments and the water column. Biological oxidation preferentially uses 12CCH4, causing the ∂13CCH4 value to increase. Carbon isotope ratios of CH4 can also allow us to differentiate oxidation and dilution processes in the water column. We calculated a Rayleigh distillation model using the equation by [Coleman et al., 1981]:	  𝛿$%𝐶' = 1000	  ×	   $+ − 1 	  × ln 𝑓 + 	  𝛿$%𝐶1. The starting stable isotope composition (∂13C0) is -40‰, and f is the fraction of the residual methane remaining (initial = 31 nmol/L) (Figure 11). The calculated oxidation curve has a kinetic isotope fractionation factor (α) of 1.002, which is in the range of expected values for oxidation processes [Whiticar and Faber, 1986]. Samples with high CH4 concentrations have low ∂13CCH4 values, while samples with low CH4 concentrations have high ∂13CCH4 values due to the enrichment of 13C during oxidation. Slight deviations from the oxidative curve are expected, as the initial concentration and stable isotope composition of the CH4 will not be identical everywhere. As the measurements follow the oxidation curve, they are not considerably altered by dilution which would cause lower CH4 concentrations to also have low ∂13CCH4 values. This affirms that oxidation is the main sink for CH4 in the water column.  Figure 11: Oxidation curve calculated using a Rayleigh distillation model with an initial concentration of 31 nmol/L, a ∂13CCH4 value of -40‰ VPDB, and an isotopic fractionation factor of 1.002. CH4 (nmol L-1 )05101520253035δ 13CCH4-42-41-40-39-38-37-36-35-34-33(‰VPDB) 33 3.2.2   CH4 in the Canada Basin and Beaufort Sea Like N2O, the CH4 distribution pattern in the Canada Basin and Beaufort Sea was similar at all of the stations. CH4 concentrations were slightly supersaturated in the PML (<5.3 nmol/L, <130% saturation), increasing to a subsurface maximum in the PSW (up to 11.5 nmol/L, ~300% saturation), and then declining through the CBIW and CBDW (Figure 13). Surface concentrations were lower than those in the PSW likely due to melt-water dilution in the polar mixed layer and air-sea flux (Section 3.3). The slight surface super-saturation could be attributed to a combination of upward diffusion from the PSW maximum and a sea ice source.  Brine channels within sea ice are frequently anaerobic, providing the necessary conditions for anoxic processes such as methanogenesis [Rysgaard et al., 2008]. In the winter, the CH4 is mixed throughout the winter mixed layer with the rejected brine, and as the ice melts in the spring, any remaining brine is released into the surface melt water layer [Damm et al., 2015a]. This process has been recently shown to cause similar surface super-saturation in the central Arctic Ocean [Damm et al., 2015a].  The PSW contained a surprisingly widespread and consistent CH4 maximum (averaging 9.1 nmol/L, 250% saturation) at a density of 25.5 σt (S= 31.8) and associated with the chlorophyll maximum (Figure 13). This peak is surprising because CH4 is normally consumed relatively quickly in oxygenated waters, and should not be produced under aerobic conditions [Reeburgh, 2007]. As discussed above, DMSP metabolism may be a source of water column CH4 production, particularly under NO3- limiting conditions [Damm et al., 2010, 2015b].  This relationship implies that DMSP may be an in-situ source of these high CH4 concentrations, providing an explanation for this ubiquitous signal.  34 Figure 12: Map and depth profiles of CH4 concentrations in the Canada Basin and Beaufort Sea  At depths below 200 m in the Canada Basin, CH4 concentrations were extremely variable, ranging from 1.0 nmol/L (27% saturation) to 24.3 nmol/L (~700% saturation), with a median value of 3.1 nmol/L (~90% saturation) below 1000 m. This under-saturation could reflect lower atmospheric CH4 concentrations at the time of ventilation (as discussed above for N2O), and/or oxidation of CH4 by methanotrophs in the deep water column. Unlike N2O, the CH4 under-saturation was not consistently observed throughout the Fluorescence0.1 0.2 0.3 0.4 0.5Depth (m)-200-160-120-80-400CH4 (nmol L-1)10 15 50CB23aFigure 13: This is a representative depth profile of CH4 and fluorescence, showing the co-occurrence of the CH4 and chlorophyll maxima. Filled black circles denote CH4 concentration and error bars represent the standard deviation of the duplicates. The dotted line is fluorescence.  35 Canada Basin, making it more difficult to attribute to a particular process. Rather, the large range of CH4 concentrations suggest the importance of ephemeral point sources. Notwithstanding the highly variable nature of CH4 concentrations, we are still able to examine the potential contribution of biological and physical processes to the observed under-saturation at certain stations. The CBDW has been isolated for ~450 years [Schlosser et al., 1997; Timmermans et al., 2003]. At the time that this water mass was last exposed to the atmosphere, the CH4 mixing ratio was ~675 ppb [Raynaud and Chappellaz, 1993]. Assuming that the temperature and salinity characteristics of this water mass have not been appreciably altered since it left the surface, the predicted equilibrium CH4 would be 1.4 nmol/L, which is 1.7 nmol/L lower than our observations. Even if some changes in temperature and salinity did occur, this would only have a small effect on our calculations. For instance, a difference in temperature of 1°C would change the solubility of CH4 by less than 0.1 nmol/L, while a salinity change of 1‰, would alter CH4 solubility by ~0.03 nmol/L. Our results thus suggest that CH4 has been added to the deep water by other, presumably biological processes. In contrast, the minimum CH4 concentration we measured was only 1.0 nmol/L, which is 0.4 nmol/L less than the atmospheric equilibrium concentration 450 years ago. This lower value, therefore, can only have been obtained through biological consumption. The lack of persistently high CH4 concentrations in the deep basin thus provides additional evidence for the microbial oxidation in these waters. In addition to very low CH4 concentrations, we observed more episodic, deeper maxima in the Canada Basin. Two of these CH4 maxima are particularly large and warrant further explanation. There was a maximum on the Beaufort shelf at 170 m (18 nmol/L), and just off the slope at 1000 m (24 nmol/L). Considering that both of these stations are deeper than the maximum depth of subsea permafrost on the Beaufort shelf (~120m), there are three possible sources that may explain these high concentrations: destabilized hydrates on the Beaufort Slope, methanogenesis in sediments, and thermogenic seeps.  Gas hydrates are stable only under specific temperatures and pressures and underlie much of the Arctic Ocean below a depth of ~300m [Ruppel, 2015]. The temperature of the water on the Beaufort Shelf has changed considerably over the last 60 years, potentially leading to a de-stabilization of gas hydrates between 300 and 550 m [Ruppel, 2015]. Indeed, “old” (low 14C) CH4 has been found in gas vents on the Beaufort shelf, implying that it is sourced from gas  36 hydrates [Paull et al., 2011]. Secondly, sedimentary methanogenesis may produce high CH4 concentrations. The maximum in the western Beaufort shelf, which is contrasted by low concentrations on the eastern shelf (mean = 4.1 nmol/L), is in agreement with the observations of Coffin et al. [2013], who measured elevated CH4 flux from the sediments to the water column of the western Beaufort Sea (near the BL stations) and low sediment-water column fluxes in the eastern Beaufort (near the MK stations). This sedimentary CH4 was primarily biogenic, implicating methanogenesis as the dominant CH4 source. Lastly, the entire Canada Basin is underlain by oil and gas deposits [Ruppel, 2015]. Gas may migrate through fractured rock and enter the sediments, and may be released through vents in the sea floor.  3.2.3    CH4 in the Canadian Arctic Archipelago Throughout the Canadian Arctic Archipelago, CH4 concentrations decreased from south (mean 7.3 nmol/L) to north (mean 3.8 nmol/L). In contrast to the other regions in our study area, we observed extremely high concentrations in surface waters of the CAA, with maximum values in excess of 420% saturation (15.3 nmol/L), and decreasing concentrations with depth (Figure 14). This surface maximum was most prominent at the southern stations (AN-314, AN-312, and VS), but was also observed at some of the northern stations (CAA-4 and CAA-7) (Figure 14). Super-saturation of CH4 in surface waters of the CAA was previously observed by Kitidis et al. [2010], who ascribed it to greater ice cover in the CAA limiting sea-air flux. Indeed, we observed high ice cover at many of the stations with elevated surface CH4 concentrations in the CAA.   37 Figure 14: Map and depth profiles of CH4 concentrations in the CAA and Baffin Bay While increased sea ice cover may contribute to the accumulation of high CH4 in CAA surface waters, ice cover, per se, may also represent a CH4 source. We found a strong negative correlation between CH4 and salinity across all regions we sampled in the CAA (Figure 15), suggesting a freshwater source of CH4 from ice and/or rivers. Rivers could supply CH4 from degrading permafrost or methanogenesis in riverine sediments to the CAA.  Sea ice, on the other hand, contains brine channels which can contain high concentrations of  CH4 [Damm et al., 2015a]. As the sea ice begins to melt, this brine drains into the surface layer of the ocean. The super-saturation we observe in the surface water at many of the stations in the CAA may be due to this partial melting of sea ice acting both as a CH4 source and a semi-permeable barrier to sea-air flux.   38  Figure 15: CH4 vs. salinity scatter plot of all samples collected in the CAA and Baffin Bay East of Lancaster Sound, in Baffin Bay, we observed a modest CH4 maximum (4.7 nmol/L, ~130% saturation) in the upper 50 m of the water column, with under-saturated concentrations below 100 m, reaching a minimum of <1 nmol/L. Our results are consistent with those obtained by Punshon et al. [2014] who observed near-equilibrium concentrations in the top ~200m of Baffin Bay, with increasing under-saturation in deeper waters. These authors attributed the depth-dependent decrease in CH4 to microbial methane consumption. In contrast with our results, Kitidis et al. [2010] observed high concentrations in the upper 150 m water column in Baffin Bay, in a depth profile taken in close proximity to our sampling station BB3. The reason for this discrepancy is unclear, but highlights the temporal variability in CH4 concentrations. 3.3   SEA-AIR CH4  AND N2O FLUXES  Sea-air fluxes of both CH4 and N2O across our full survey region were low compared to previous estimates from other Arctic regions, but similar to those from the global ocean (Table 2, Table 3) [Rhee et al., 2009]. Overall, our study area was a source of CH4 with sea-air flux averaging 1.3 µmol m-2 d-1 (range = -0.4 to 4.9 µmol m-2 d-1), and a minor sink for N2O, with an average sea-air flux of -0.52 µmol m-2 d-1 (range = -4.1 to 2.5 µmol m-2 d-1). Salinity (‰)22 24 26 28 30 32 34 36CH4 (nmol L-1)0246810121416R2 = 54.9%p<0.0001 39 These findings are comparable to typical sea-air fluxes from the world ocean, which range from 0.2-1.5 µmol m-2 d-1 and 0.3-0.6 µmol m-2 d-1, for CH4 and N2O, respectively [Rhee et al., 2009]. At station CB28aa in the eastern Beaufort Sea, we observed high sea-air fluxes of both CH4 and N2O.  At this station, we observed upwelling of deep water onto the shelf (based on its high density) implying that upwelling is a source of these gases to the atmosphere. This observation is in good agreement with previous studies showing the importance of coastal upwelling as a source of elevated CH4 and N2O sea-air fluxes in various marine systems [eg., Nevison et al., 2004; Frame et al., 2014; Capelle and Tortell, 2016]. Table 2: Sea-air flux of N2O  Although our N2O sea-air fluxes are relatively close to previous studies (Table 2), our CH4 fluxes are much lower than those measured on Eurasian shelves, in CH4 “hotspots”, and in polynya waters (Table 3). There are several reasons for this. Instead of taking the instantaneous ship wind speed, as is often done for air-sea flux calculations, we used a weighted mean over the 60 days prior to sampling. This approach incorporated variability in wind-speed, and thus produces piston velocity estimates that are more reflective of the average conditions experienced over the time of mixed layer ventilation.  The more recent wind speeds are more heavily weighted, causing the wind speeds within the 7 days prior to sampling to have the most influence on the piston velocity (kw). These derived weighted piston velocities are often lower than those based on instantaneous or daily mean wind speed data.  Indeed, we obtained significantly higher (by up to 190%) sea-air fluxes when using mean daily wind speeds. An additional explanation for our lower derived CH4 sea-air fluxes relates to the strong spatial variability in CH4 concentrations. Unlike many Region Sea-air flux (µmol N2O m-2 d-1) Reference Bering Sea 2 ± 1 -0.5 ± 0.5 [Hirota et al., 2009] This study Western Chukchi Sea 12 ± 9 [Hirota et al., 2009] Eastern Chukchi Sea -0.4 ± 0.7 This study Canada Basin -0.9 ± 1.4 This study CAA 5 ± 8 -0.2 ± 1.1 [Gagné, 2015] This study  40 previous studies, we did not specifically sample any areas considered “hotspots” for CH4, so our estimates should be considered a conservative background level. Understanding the significance of CH4 hotspots to Arctic Ocean CH4 fluxes requires careful sampling that avoids statistical bias in the over-representation of either high or low CH4 waters. Carefully designed sampling programs, coupled with an objective analysis will provide a better perspective on the spatial distribution of these “hotspots”, as well as a more complete picture of the overall distribution of CH4 concentrations. Although sampling biases and different calculation approaches may explain the somewhat lower sea-air fluxes we observed relative to previous studies, there are also several hydrographic properties that would act to limit sea-air gas exchange. In general, low sea-air fluxes of CH4 and N2O can be attributed to the dilution of surface gas concentrations by melt-water, as previously suggested by Kitidis et al. [2010]. Indeed, fresh water contained low surface concentrations of both gases (Figure 2, Figure 3), resulting in low CH4 and N2O fluxes in the upper layer that exchanges with the atmosphere (Table 2, Table 3). As we did not have measurements within the mixed layer for stations that were partially covered with sea-ice, we did not calculate fluxes for these locations. Table 3: Sea-air flux of CH4 Region Sea-air flux (µmol CH4 m-2 d-1) Reference Bering Sea 0.2 ± 0.4 This study Eastern Chukchi Sea 1.9 ± 1.4 This study Western Chukchi Sea 5 – 57 [Savvichev et al., 2007] Eastern Siberian Arctic (area-weighted mean flux, summer) ~7 - 17 [Shakhova and Semiletov, 2007] East-Siberian Arctic Sea (CH4 hotspots, summer) 45 - 95 [Shakhova and Semiletov, 2007] Barents Sea ~5 [Lammers et al., 1995] Canada Basin 1.3 ± 1.1 This study CAA 1.2 ± 1.1 This study Davis Strait ~1.6 [Punshon et al., 2014] Storfjorden polynya 26 - 104 [Damm et al., 2007]  Based on the average sea-air flux for each sub-region and its area (Table 4), we calculated the annual emission of CH4 and N2O for the entire study region. We estimated  41 an annual flux of CH4 on the order of 0.014 ± 0.013 Tg CH4 yr-1, equating to approximately 0.14 % of the yearly global oceanic CH4 emissions, which are thought to be ~10 Tg yr-1 [Reeburgh, 2007]. This is a conservative estimate, as it does not take into account CH4 accumulation under sea ice and subsequent release after sea ice break up, storm events (that increase the depth of the mixed layer), or “hotspots” that may be significant local sources of CH4. It is important to note that the Eastern Siberian Seas have been estimated to emit 1 – 8 Tg CH4 yr-1 alone [Shakhova et al., 2005, 2010a], so our sea-air fluxes represent a minimum for the Arctic Ocean. We found that the North American Arctic Ocean was a minor sink of N2O, with an integrated sea air flux on the order of ~-0.01 ± 0.023 Tg N yr-1. This low balance can be attributed to freshwater input in the summer which creates a low dissolved gas layer at the surface of the Arctic Ocean [Macdonald et al., 2004]. Our Chukchi Sea flux is much lower than previous estimates by Hirota et al. [2009], who determined the flux in the entire Chukchi Sea to be ~0.07 Tg N yr-1 (Table 2). However, this previous calculation was based on much fewer measurements, of which the majority were in the western region. Thus, we consider our estimate to be more representative for the eastern Chukchi Sea. Table 4: Regional annual fluxes of CH4 and N2O Region Mean sea-air flux  (µmol m-2 d-1) Area (km2) Annual flux Northern Bering Sea -0.5 ± 0.5 N2O  0.2 ± 0.4 CH4  320,000  -1.7 ± 1.5 Gg N 0.3 ± 0.7 Gg CH4 Eastern Chukchi Sea -0.4 ± 0.7 N2O  1.9 ± 1.4 CH4 210,000 -0.9 ± 1.4 Gg N  2.3 ± 1.7 Gg CH4  Canada Basin -0.9 ± 1.4 N2O  1.3 ± 1.1 CH4  700,000 -6.1 ± 10.2 Gg N  5.5 ± 4.6 Gg CH4  CAA -0.2 ± 1.1 N2O 1.2 ± 1.1 CH4 840,000 -0.9 ± 9.6 Gg N  5.7 ± 5.5 Gg CH4    Total flux: -0.0097 ± 0.023 Tg N  0.014 ± 0.013 Tg CH4    42 4    Conclusions 4.1   MAIN FINDINGS This study covers a larger sampling region than any previous depth-resolved study of either CH4 or N2O in the Arctic Ocean. This synoptic perspective enables a pan-arctic view of the distributions of these gases, and reveals several important patterns. The highest concentrations of both gases were measured in the bottom waters of the shallow Chukchi Sea, where active microbial metabolism under low O2 conditions leads to high rates of net CH4 and N2O production. The high N2O concentrations that result from the denitrification and nitrification in the sediments on the shelf, are transported through the Canada Basin until they are attenuated in the CAA. In contrast, the CH4 is rapidly oxidized in the water column so that the high CH4 signature has a much more limited spatial extent. Both large super-saturation and under-saturation of CH4 were measured, likely owing to the diversity and ephemerality of its sources, and the high capacity for CH4 consumption in oxygenated waters. Indeed, the heavy ∂13CCH4 signatures of low CH4 concentrations were likely a result of consumption within the water column, indicating that this is a key sink of CH4. Conversely, sources of CH4 in the Arctic Ocean may result from sedimentary methanogenesis, decomposing permafrost, dissociating gas hydrates, migration from geologic gas deposits, river input, methanogenesis in sea ice, and in situ water column production. Additionally, sea ice formation and melting were important processes in the cycling of these climate-active gases. Convective overturning due to brine rejection stirred up the sediments, likely adding CH4 and N2O to the water column where seafloor production of these gases is high. Production within sea ice was also likely a source, and may have caused the super-saturation of CH4 areas where sea ice was only partially melted. Although we measured elevated concentrations of both of CH4 and N2O in the water  43 column, sea-air fluxes of these gases were relatively low compared to previous studies in the Arctic Ocean [e.g., Hirota et al., 2009; Shakhova et al., 2010]. The freshwater layer at the surface acted to dilute the concentrations of the gases within the mixed layer, limiting the associated sea-air fluxes.   This study adds a major new set of observations to the currently limited data set of Arctic Ocean CH4 and N2O, with the first depth profiles of CH4 in the Canada Basin and Canadian Arctic Archipelago (CAA), and N2O depth profiles in the Beaufort Sea and much of the CAA. Our large-scale data set, combined with other hydrographic and oceanographic measurements, allowed us to examine likely processes driving the cycling of CH4 and N2O in the North American Arctic Ocean. We show that the Pacific Winter Water mass (PWW) is characterized by the highest N2O concentrations in the North American Arctic, and this signature persists as far east as the CAA. We also examined the importance of permafrost in relation to the water column concentrations of CH4 by comparing the sea-air fluxes in the Eastern Siberian Arctic Sea with those from the permafrost-free Chukchi Sea. We studied the role of sea ice in the biogeochemical cycling of CH4 and N2O. We documented the subsurface CH4 maximum in the Canada Basin as well as the surface super-saturation, but under-saturation at depth at stations partially covered with sea-ice in the CAA.  A major contribution of this thesis is the provision of baseline CH4 and N2O concentration data for the Arctic.  These baseline data will provide a benchmark for comparison against future studies as the Arctic environment undergoes rapid changes.  Based on the results we obtained, we are able to speculate on how CH4 and N2O accumulation in Arctic waters may be influenced by climate-driven environmental change. The outcome will mainly depend on changes in carbon export and organic matter availability in shallow sediments, the reduction of the sea ice barrier to air-sea exchange, and the stability of shallow gas hydrate deposits on continental shelves. 4.2   CAVEATS AND LIMITATIONS As this study was conducted on three separate cruises in different regions of the North American Arctic, we did not obtain the same resolution in all areas. The Beaufort Sea and CAA did not have the same density of stations as the northern Bering Sea and Chukchi Sea,  44 limiting the conclusions we could draw from these regions. We also did not have as many ∂13CCH4 isotope samples in the Canada Basin and Beaufort Sea or the CAA as we did in the Bering and Chukchi Seas, so we were unable to determine the CH4 sources east of the Chukchi Sea. As we only conducted sampling during the summer, our temporal resolution is also limited and we are unable to resolve seasonal changes in CH4 and N2O concentrations. 4.3   FUTURE OUTLOOK This study reveals interesting new questions to guide future research. Additional CH4 studies should focus characterizing CH4 distributions in areas with subsea permafrost. The Beaufort Sea and Banks Island are the only regions in the North American Arctic Ocean with significant subsea permafrost deposits [Ruppel, 2015]. However, the Beaufort Sea has not been analyzed for CH4 in any comprehensive study since the early 1990’s [Macdonald, 1976; Kvenvolden et al., 1993], and no measurements have been made in the vicinity of Banks Island.  Additional samples for isotopic analysis, particularly 14C, would also be especially useful to differentiate between CH4 sourced from the methanogenesis of recent carbon, or CH4 derived from old carbon sources such as degrading permafrost and gas hydrates.  Future studies should also investigate the source of the ubiquitous subsurface maximum in the Canada Basin. This feature is unusual because of the low primary productivity, yet higher CH4 concentrations than typically observed in open ocean subsurface maxima (~3-5 nmol/L) [Reeburgh, 2007]. Potential processes leading to this sub-surface maximum could be investigated using  stable isotopes (∂13CCH4 and ∂DCH4), CH4 production rate measurements, and DMSP-addition experiments.  Future N2O research should focus on following the PWW as it transits through the northern CAA. It remains unclear whether the N2O from the maximum is ventilated to the atmosphere or is diluted by river input and meltwater. Sampling with high spatial resolution in the northern archipelago may elucidate the specific processes that cause the decrease in N2Oconcentrations as the water moves eastward. Additionally, N2O studies should focus on the Eurasian shelves. N2O distributions have not been studied outside of the North American Arctic Ocean, although the low oxygen concentrations and decomposing permafrost on these marginal shelves may produce significant quantities of  45 N2O. Finally, the relationship between sea-ice cycling and CH4 / N2O dynamics is still poorly understood. In our study, fresh water acted to limit sea-air flux (due to dilution of gas concentrations), while we observed high gas concentrations in the surface water under marginal sea-ice conditions. It remains unclear what happens to the gas distribution as sea-ice breaks up in the spring and re-forms in the fall. High temporal resolution sampling (depth profiles, sea ice, and brine) during these critical times could provide an explanation for the high super-saturation observed under sea ice, relative to the underlying water column. It may also help us understand the sources of CH4 and N2O to the water column, and the seasonal variability in their sea-air flux.   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