thermokarst lakes on the arctic coastal plain of alaska: geomorphic controls on bathymetry

13
Thermokarst Lakes on the Arctic Coastal Plain of Alaska: Geomorphic Controls on Bathymetry Kenneth M. Hinkel, 1 * Yongwei Sheng, 2 John D. Lenters, 3 Evan A. Lyons, 2 Richard A. Beck, 1 Wendy R. Eisner 1 and Jida Wang 2 1 Department of Geography, University of Cincinnati, Cincinnati, OH, USA 2 Department of Geography, University of California, Los Angeles, CA, USA 3 School of Natural Resources, University of Nebraska, Lincoln, NE, USA ABSTRACT Detailed bathymetric data were collected for 28 thermokarst lakes across the Arctic Coastal Plain (ACP) of northern Alaska from areas with distinctly different surcial sediments and topography. Lakes found in the low-relief coastal area have developed in marine silts that are ice-rich in the upper 610 m. The lakes tend to be shallow (~ 2 m), of uniform depth and lack prominent littoral shelves. Further inland on the ACP, lakes have formed in relatively ice-poor aeolian sand deposits. In this hilly terrain, average lake depth is less (~ 1 m) despite deeper (35 m) central pools. This bathymetry reects the inuence of broad, shallow littoral shelves where sand, eroded from bluffs at the lake margin, is deposited concurrently with deep penetration of the talik beneath the basin centre. Lakes in the ACP-Arctic Foothills transition zone to the south have developed in loess uplands. These yedoma deposits are extremely ice-rich, and residual lakes found inside old lake basins (alases) are generally 24 m deep, reecting continued talik development and ground subsidence following drainage of the original lake. However, where the expanding lake encroaches on the anks of the upland at actively eroding bluffs, near-shore pools develop that can be 69 m deep. It appears that thawing of ice-rich permafrost during lake expansion causes ground subsidence and formation of deep pools above ablating ice wedges. These data suggest that thermokarst lake morphometry largely depends on the characteristics of the substrate beneath the lake and the availability of sediments eroded at the lake margin. Copyright © 2012 John Wiley & Sons, Ltd. KEY WORDS: Alaska; thermokarst lakes; limnology; permafrost; bathymetry INTRODUCTION Tens of thousands of thermokarst lakes and ponds are found in the continuous permafrost region of northern Alaska. Pooling of water in summer causes reduced surface albedo, leading to localised warming and degradation of the permafrost immediately beneath the pool. If the substrate comprises frozen sediments that contain excess (non-pore) ground ice, then thawing permafrost settles and the ground surface consequently subsides to yield a deeper pool. Over time, pools can deepen, enlarge and coalesce to form thermokarst lakes that expand in area and volume, and can persist for thousands of years. The evidence to date suggests that thermokarst lakes are increasing in size and number due to enhanced regional climate warming (Smith et al., 2005; Riordan et al., 2006), and this will likely have a notable impact on lake dynamics and permafrost stability (Jorgenson et al., 2006). Lake basin deepening by subsidence of thawing permafrost increases the volume of the unfrozen water reservoir beneath the ice in winter (Brown and Duguay, 2010). Given the high heat capacity of water, this represents a substantial net increase in the amount of energy available for warming and thawing basal sediments (Plug and West, 2009). Thawing permafrost is associated with enhanced microbial activity and mobilisation of sequestered soil organic carbon, with the result that thermokarst lakes are an important source of CO 2 and biogenic CH 4 to the atmosphere (Kling et al., 1991; Zimov et al., 1997; Walter et al., 2006, 2007). Assessing the likely impact of climate warming on thermokarst lakes is difcult, however, because little baseline limnological information has been collected from a * Correspondence to: Kenneth Hinkel, Department of Geography, University of Cincinnati, Cincinnati, OH 45221-0131, USA. E-mail: [email protected] PERMAFROST AND PERIGLACIAL PROCESSES Permafrost and Periglac. Process. (2012) Published online in Wiley Online Library (wileyonlinelibrary.com) DOI: 10.1002/ppp.1744 Copyright © 2012 John Wiley & Sons, Ltd. Received 5 September 2011 Revised 21 June 2012 Accepted 6 July 2012

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PERMAFROST AND PERIGLACIAL PROCESSESPermafrost and Periglac. Process. (2012)Published online in Wiley Online Library(wileyonlinelibrary.com) DOI: 10.1002/ppp.1744

Thermokarst Lakes on the Arctic Coastal Plain of Alaska: Geomorphic Controlson Bathymetry

Kenneth M. Hinkel,1* Yongwei Sheng,2 John D. Lenters,3 Evan A. Lyons,2 Richard A. Beck,1 Wendy R. Eisner1 and Jida Wang2

1 Department of Geography, University of Cincinnati, Cincinnati, OH, USA2 Department of Geography, University of California, Los Angeles, CA, USA3 School of Natural Resources, University of Nebraska, Lincoln, NE, USA

* CoUnivE-ma

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ABSTRACT

Detailed bathymetric data were collected for 28 thermokarst lakes across the Arctic Coastal Plain (ACP) of northernAlaska from areas with distinctly different surficial sediments and topography. Lakes found in the low-relief coastalarea have developed in marine silts that are ice-rich in the upper 6–10m. The lakes tend to be shallow (~ 2m), ofuniform depth and lack prominent littoral shelves. Further inland on the ACP, lakes have formed in relatively ice-pooraeolian sand deposits. In this hilly terrain, average lake depth is less (~ 1m) despite deeper (3–5m) central pools. Thisbathymetry reflects the influence of broad, shallow littoral shelves where sand, eroded from bluffs at the lake margin, isdeposited concurrently with deep penetration of the talik beneath the basin centre. Lakes in the ACP-Arctic Foothillstransition zone to the south have developed in loess uplands. These yedoma deposits are extremely ice-rich, and residuallakes found inside old lake basins (alases) are generally 2–4m deep, reflecting continued talik development and groundsubsidence following drainage of the original lake. However, where the expanding lake encroaches on the flanksof the upland at actively eroding bluffs, near-shore pools develop that can be 6–9m deep. It appears that thawingof ice-rich permafrost during lake expansion causes ground subsidence and formation of deep pools above ablatingice wedges. These data suggest that thermokarst lake morphometry largely depends on the characteristics of thesubstrate beneath the lake and the availability of sediments eroded at the lake margin. Copyright © 2012 JohnWiley& Sons, Ltd.

KEY WORDS: Alaska; thermokarst lakes; limnology; permafrost; bathymetry

INTRODUCTION

Tens of thousands of thermokarst lakes and ponds are foundin the continuous permafrost region of northern Alaska.Pooling of water in summer causes reduced surface albedo,leading to localised warming and degradation of the permafrostimmediately beneath the pool. If the substrate comprisesfrozen sediments that contain excess (non-pore) ground ice,then thawing permafrost settles and the ground surfaceconsequently subsides to yield a deeper pool. Over time,pools can deepen, enlarge and coalesce to form thermokarstlakes that expand in area and volume, and can persist forthousands of years.

rrespondence to: Kenneth Hinkel, Department of Geography,ersity of Cincinnati, Cincinnati, OH 45221-0131, USA.il: [email protected]

right © 2012 John Wiley & Sons, Ltd.

The evidence to date suggests that thermokarst lakes areincreasing in size and number due to enhanced regionalclimate warming (Smith et al., 2005; Riordan et al., 2006),and this will likely have a notable impact on lake dynamicsand permafrost stability (Jorgenson et al., 2006). Lake basindeepening by subsidence of thawing permafrost increasesthe volume of the unfrozen water reservoir beneath the icein winter (Brown and Duguay, 2010). Given the high heatcapacity of water, this represents a substantial net increasein the amount of energy available for warming and thawingbasal sediments (Plug and West, 2009). Thawing permafrostis associated with enhancedmicrobial activity andmobilisationof sequestered soil organic carbon, with the result thatthermokarst lakes are an important source of CO2 andbiogenic CH4 to the atmosphere (Kling et al., 1991; Zimovet al., 1997; Walter et al., 2006, 2007).

Assessing the likely impact of climate warming onthermokarst lakes is difficult, however, because little baselinelimnological information has been collected from a

Received 5 September 2011Revised 21 June 2012Accepted 6 July 2012

K. M. Hinkel et al.

regionally representative sample of lakes. Remotelysensed imaging systems have been used to derive general-ised lake bathymetry, but are most effective for shallow(< 2m) lakes (Mellor, 1994; Morris et al., 1995; Jeffrieset al., 1996; Kozlenko and Jeffries, 2000). Numericalmodels based on conductive heat transfer have been usedto simulate talik development and basin morphology(Pelletier, 2005; West and Plug, 2008), but have inherentlimitations (Plug and West, 2009). This study reports ona project designed to obtain detailed bathymetric measure-ments for a group of lakes along a 150-km long north-southtransect across the Arctic Coastal Plain (ACP), andexamines the relation between basin morphometry andregional geomorphology. The purpose is to collect baselinebathymetric data from lakes developed in varyinglandscape units to identify fundamental controls on lakebasin evolution.

Figure 1 Shaded relief map of northern Alaska showing topography, physiografigure is available in colour online at w

Copyright © 2012 John Wiley & Sons, Ltd.

STUDY AREA AND BACKGROUND

Geologic Conditions

TheACP of northern Alaskawas formed largely by sedimentseroded from the east-west-trending Brooks Range and trans-ported northward toward the Arctic Ocean, where they wereextensively reworked by marine, fluvial and lacustrineprocesses. The area is underlain by continuous permafrost thatis 300–600m deep, and has low elevation and relief, orientedthermokarst lakes and drained lake basins, and north-flowingmeandering or braided streams that dissect the region (seeFigure 1). Further inland towards the Brooks Range is theArctic Foothills physiographic province (Wahrhaftig, 1965);the boundary between the ACP and Foothills is detectableat 120–200m above sea level (asl). The Arctic Foothillsexperienced piedmont glaciation during the Last Glacial

phic regions, villages and sampled lakes including focus lakes A–C. Thisileyonlinelibrary.com/journal/ppp

Permafrost and Periglac. Process., (2012)

Table

1Generalised

stratig

raphic

sequence

andcharacteristicsforthethreegeom

orphic

regionsin

thestudyarea.

Outer

Coastal

Plain

InnerCoastal

Plain

ACP-A

rctic

Foothillstransitio

nzone

Barrow

unitof

theGubik

Formation:

Quaternary-agesediments,mostly

marine,but

also

fluvialandlacustrine.Silt

togravel,marine

andlacustrine

siltcommon

atthesurface.

Com

plex

historywith

migratin

gshorelines

and

emerging

beaches,bars,andlagoonsextensively

reworkedby

stream

andthermokarst

lake

processes.Generally

abundant

organicmatter

andice,especially

icewedgesto

6–10

m.Grades

laterally

into

theMeade

River

unit

Pleistocene

sand

seaas

thin

layeror

dunesup

to35

mthick;

generally

ice-poor

Massive

aeoliansandysilts

andfine

sand

with

silt

cap,

very

ice-rich

with

deep

(>10

m)syngenetic

icewedges.‘Silt

belt’

ofCarter(1988)

identifi

edby

Kanevskiy

etal.(2011)as

yedoma.Deposited

asaplateau20

–30m

thick,

thickerin

valleys.

Subsequentstream

incision

andmassmovem

ent.

Contemporaneouswith

Pleistocene

sand

sea

Meade

River

unitof

theGubik

Formation.

Generally

well-sorted

quartzsand

ofmarineorigin,

butalso

fluvialsandsandgravels.Overliesolder

SkullCliffunitor

unconformably

onbedrock,

especially

tothewest.

Sandseaandupland

siltfacies

originally

included

inthisunit

Meade

River

unitof

theGubik

Formation

Cretaceous-Tertiary

sedimentary

bedrock

Cretaceous-Tertiary

sedimentary

bedrock

Cretaceous-Tertiary

sedimentary

bedrock

Black

(1964);B

rown(1968);B

rownetal.(2001

);Sellm

annet

al.(1975)

Black

(1964);Brownet

al.(2001);

Carter(1981,

1988);Williamset

al.(1978);

Williams(1983)

Brownet

al.(2001);Carter(1988);

Kanevskiy

etal.(2011);Law

son(1983);

WilliamsandYeend

(1978);Williams(1983)

ACP=Arctic

Coastal

Plain.

Arctic Lakes Bathymetry

period, and are characterised by rolling topography, glacialdeposits and aeolian silt extending up to about 1500m asl.Lakes here tend to be less common, are not oriented, and mostare either moraine dammed or kettle lakes developed inresponse to glacial processes (Kling et al., 1992).The ACP experienced seven to eight marine transgressions

during the late Cenozoic (Dinter et al., 1990; Brigham-Gretteand Carter, 1992; Brigham-Grette and Hopkins, 1995;Brigham-Grette, 2001), and the more recent events have leftevidence of higher sea levels. Hopkins (1973, 1982) andothers (e.g. Péwé, 1975; Carter, 1993) have identified ancientshorelines and wave-cut scarps that can be discontinuouslytraced across the ACP. Of relevance to this study is a promi-nent ancient shoreline at 23–29m asl that is associatedwith a pre-Illinoian interglacial high stand predating about175 ka (Lewellen, 1972; Sellmann et al., 1975; Hopkins,1982). The effects of prolonged coastal submergence arereflected in the surficial deposits and topography, since theancient shoreline forms a boundary between the flat-lyingOuter Coastal Plain (OCP) to the north composed of surficialmarine silt and sand and the rolling topography of the InnerCoastal Plain (ICP) of surface aeolian sand and silt to thesouth (Williams et al., 1978; Williams, 1983).Surface deposits in the western ACP are members of the

Gubik Formation, unconsolidated Quaternary-age sedimentsof dominantly marine origin that lie unconformably abovenearly horizontal sedimentary bedrock of Cretaceous orTertiary age (see Table 1). On the OCP, the Barrow unit ofthe Gubik Formation is a complex mix of poorly sorted towell-sorted clay, silt, sand and gravel some 10–25m thick thatis found within 30 km of the modern coast and within 15m ofsea level (Black, 1964). Silt tends to dominate near the surfaceowing to the effects of marine, fluvial and thermokarst lakeprocesses, and near-surface silt deposits are especially ice-rich,often exceeding 80 per cent ice by volume (Brown, 1968;Sellmann et al., 1975). In addition to interstitial (pore) ice,the unit contains ice veins, lenses, networks and ice wedges.Ice wedges tend to be broad and shallow, such that the excessice content is greatly reduced below a depth of 6–10m (Carsonand Hussey, 1962; Lewellen, 1972; Sellmann et al., 1975).South of the ancient shoreline on the ICP is a large aeolian

sand sea that developed during the cold and windy latePleistocene (Williams et al., 1978; Carter, 1981, 1993;Williams, 1983). Fine to medium sand forms a mantle thatcan exceed 15m in thickness. Although pore and wedge iceexist, the unit lacks large volumes of excess ice and is consid-ered thaw stable. Beneath the Pleistocene sand sea on the ICPis the Meade River unit of the Gubik Formation, which iscontemporaneous with the Barrow unit but is somewhatthicker (up to 60m). It generally contains coarser sedimentsconsistent with the more proximal location relative to thesediment source. Composed of quartz sand and silt, this unitreflects the complex late Cenozoic history with migratingshorelines and mixed marine and fluvial sediments. Creta-ceous bedrock crops out near the beds of larger streams suchas the Meade River, and the unit appears to thicken eastwardsbeneath the surficial aeolian sands (Black, 1964).

Copyright © 2012 John Wiley & Sons, Ltd. Permafrost and Periglac. Process., (2012)

K. M. Hinkel et al.

Near the southern boundary of the ACP are thickaccumulations of silt and fine sand deposited subaeriallyduring the late Wisconsinan, and contemporaneously withthe Pleistocene sand sea (Carter, 1981, 1988). Originallyforming as a broad plateau, the ice-rich upland silt depositswere subsequently dissected by stream incision and affectedby mass movement processes (Williams and Yeend, 1978),and now stand as hills or ridges some 20–50m higher thanthe surrounding terrain. These deposits are part of theupland silt belt discussed by Carter (1988) that is found justsouth of the marine, fluvial and aeolian deposits of theadjacent ICP, and form a 30-km wide transition zonebetween the ICP and the Arctic Foothills. More recently,they have been identified as yedoma deposits by Kanevskiyet al. (2011, Figure 9). Yedoma is well known from Siberia(Popov, 1953; Romanovskii, 1993; Grosse et al., 2007),Canada (Fraser and Burn, 1997) and northern Alaska(Carter, 1988; Shur et al., 2004; Kanevskiy et al., 2011),and develops where sedimentation of fine material is fairlycontinuous such as deltas, flood plains and loess belts(Morgenstern et al., 2011). As the ground surface aggradesupward, the permafrost table migrates upward as well(syngenetic permafrost), creating the diagnostic ataxitic(suspended) cryostructures. Soil organic matter and plantrootlets are often incorporated into the aggrading permafrost,and syngenetic ice wedges are common. A description of thedeposits near the study area indicates that massive tabular iceand deep syngenetic ice wedges are common (Williams andYeend, 1978; Carter, 1988; Lawson, 1988), and the excessice content may exceed 70 per cent. Deep (> 20m) drainedthermokarst lake basins (alases) are known from the region(Williams and Yeend, 1978; Wang et al., 2012).

Thermokarst Lakes

Developed atop the permafrost are thousands of lakes andponds that are frozen for 8–9months of the year withmaximum ice thickness of 1–2m. Analysis of satellite imagery(Sellmann et al., 1975; Hinkel et al., 2003, 2005; Frohn et al.,2005; Wang et al., 2012) reveals that about 20 per cent of theACP is covered by lakes, while 26 per cent of the surface ischaracterised by topographic basins that are remnants of lakesthat have drained. Previous work (Hinkel et al., 2003, 2005;Frohn et al., 2005; Wang et al., 2012) has demonstrated thatthe concentration and size of thermokarst lakes decreasetoward the interior; there are more and larger lakes on theOCP than on the ICP. When including innumerable ponds,wetlands and streams, it is clear that water bodies are thedominant element on the landscape, and that lacustrineand fluvial processes are currently the primary agents ofgeomorphic work.Lake formation apparently began with widespread

permafrost degradation in the late Wisconsinan, with manylakes developing during the warmer early Holocene around10 ka ago (Ritchie et al., 1983; Hopkins and Kidd, 1988;Rampton, 1988; Mann et al., 2002; Kaufman et al., 2004;Walter et al., 2006, 2007). Lake development is associated

Copyright © 2012 John Wiley & Sons, Ltd.

with local or regional deepening of the active layer andthermokarst activity associated with the upper, ice-rich layerof permafrost (Burn, 1997; Côté and Burn, 2002; French,2007; Jorgenson and Shur, 2007; Plug and West, 2009).Once the lake is deeper than the thickness of the winterice cover (1–2m), an unfrozen reservoir of water persiststhroughout the year and maintains a net positive heat flowinto the basal sediments (Brewer, 1958; Lachenbruchet al., 1962; Smith, 1976). This occurs when the long-termmean lake-bottom temperature exceeds 0�C (Ling andZhang, 2003), making the lake bed some 10�C warmer thanthe surrounding tundra surface in the study area. Thus, lakescause a fundamental disruption in the thermal stability ofnear-surface permafrost by altering the long-term groundsurface temperature.

Over time, the talik penetrates downward into thepermafrost. Thaw subsidence increases the volume of unfrozenwater, enhances the energy content of the reservoir of waterwith its high specific heat capacity and promotes further thawin a positive feedback system. The depth of thaw frontpenetration is determined by the mean annual water tempera-ture at the bed in the central basin, the thermal characteristicsand ice content of the underlying sediments, the dimensionsof the central basin, the history of the lake and time (Burn,2002; Ling and Zhang, 2003, 2004; West and Plug, 2008;Plug and West, 2009).

At the lake surface, winds create currents and generate atwo-cell circulation pattern that causes thermomechanicalbank erosion and areal expansion, resulting in asymmetricalelliptical orientation (Livingstone, 1954; Rex, 1961; Carsonand Hussey, 1962; Mackay, 1963; Carson, 1968, 2001; Arpet al., 2011). Preferential bank erosion is concentrated inzones oriented 50� to the wave approach; this promotes bankerosion at the northern and southern ends of the lakes on theACP, and consequent lake elongation and orientation.Concurrent sediment deposition often occurs on the easternand western lake margins in the form of shallow littoralshelves. Sediments tend to be sorted during lake expansion,with the coarser faction (fine to medium sand) preferentiallydeposited as massive or layered sheets on the progradinglittoral shelves, while the finer faction (silts) and organicmatter tend to accumulate in the deeper central basin (Murton,1996). Because lake ice freezes to the bed on the shallowlittoral shelves, there is no talik beneath these features (Burn,2005). The active layer here is typically 0.6 to 0.8m deep,with little excess ice in the underlying frozen sand.

METHODOLOGY

A 3-year project (2008–10) was undertaken, with most ofthe fieldwork performed in summer. In year 1, lakes nearthe coastal village of Barrow (100 series) were studied(Figure 1). Located on the low-relief OCP, lakes here aredeveloped in extremely ice-rich marine silts and sands. Inyear 2, the project moved inland about 100 km and stagedout of the village of Atqasuk on the ICP. Lakes near here

Permafrost and Periglac. Process., (2012)

Arctic Lakes Bathymetry

(200 series) have formed in the Pleistocene sand sea. Therelief is somewhat greater, with elevations ranging from15 to 45m asl. In year 3 (300 series), the study areawas 65 km further south in an area informally known as‘Reindeer Camp’, situated in the silt belt of the ACP-ArcticFoothills transition zone. Here, marine, fluvial and aeoliansediments of the ACP grade into loess-covered hills, andlake density decreases. About 15 potential lakes for eachstudy area were selected prior to the field season to obtaina range of lake sizes. However, since lakes were ultimatelyaccessed by float plane, only larger and deeper lakes couldbe visited. Sensors were deployed in each lake to monitorsummer water temperature and water depth, and the resultsare presented in the companion paper (Hinkel et al., 2012).In addition, one ‘focus’ lake was selected from each of thethree regions for detailed year-round energy and waterbalance studies.Detailed bathymetric data were collected from lakes in

each of the study areas using a GPS-enabled sonar unit. Lakebathymetric surveys were conducted in late June (if the lakewas completely free of ice) or August. A Lowrance/EagleSeaCharter 502cDF iGPS sonar unit was attached to thetransom of a 2.7-m (90) inflatable boat. Geographic locationand water depth were recorded each second along a transect.Beginning just offshore, the first transect paralleled the shore-line for an entire circuit around the lake. The next transect wasshifted 50–100m towards the lake centre and ran parallel tothe initial transect. This process was continued until the entirelake was surveyed using these nested transects; an example isshown in Figure 2a. This procedure was modified whenspecific features were encountered. For example, in order tolocate the shelf edge accurately and to assess gradients, theboat operator would manoeuvre back and forth across theshelf edge. Similarly, when deep pools were encountered,additional transects were used to reduce the spacing andimprove the resolution for detailed mapping. Since theoutboard motor could not operate in water depths of less than0.3m, these areas were avoided.The accuracy of the depthmeasurements depended on local

conditions. The primary uncontrolled factors that degradedaccuracy were the presence of thick aquatic vegetation atdepth (rare) and wave height. Many days were windy, andthe resulting waves caused the small craft to pitch, whichchanged the angle of the transducer relative to the lakebottom. The effect is minor except in deeper water (> 3m),where an estimated error of � 0.2m was possible.The transect data were converted to an Excel time series

within the SonarViewer 1.2.3 software provided by Lowrance,Inc. Individual charts were merged to create a file containingthe latitude, longitude and water depth for every measuredpoint on the lake. The shoreline of each lake was delineatedfrom colour-infrared (CIR) aerial photographs with a 2.5-mhorizontal resolution, or an Interferometric synthetic apertureradar (IfSAR)-derived digital elevation model (DEM) with aspatial resolution of 5.0m and vertical accuracy of � 1.0mor better (Intermap Technologies, Inc., 2002; Manley et al.,2005). Locations along the shore were assigned a depth value

Copyright © 2012 John Wiley & Sons, Ltd.

of 0.0m and appended to the Excel file, which was importedinto ArcGIS as a shape file (WGS84 datum) for management,analysis and mapping. A natural neighbour algorithm wasemployed to interpolate the bathymetric surface with anominal cell spacing of 5m. If the resulting raster extendedbeyond the actual lake shoreline, the shoreline polygon shapefile was used as a mask to clip the raster. The resultingbathymetric raster was classified using an equal intervalscheme andmapped as a coloured isarithmic surface. Two typ-ical examples from each of the three study areas are shown inFigure 2, accompanied by an orthophoto (http://eros.usgs.gov/#/Find_Data/Products_and_Data_Available/Aerial_Products)and elevation profile for each lake.

Tomonitor the change in lake water level over the summer,a HOBOwater-level logger (U20-001-01, precision of 2mm)or Solinst unit (precision of 1mm) was deployed in the yearthat the lake was surveyed. The data (not presented) generallyshow higher water levels after ice-off in June, and a gradualreduction in water levels by 5–10 cm over several weeks.After mid-July, water levels tended to stabilise for the re-mainder of the summer, with some fluctuations caused byprecipitation events. These seasonal variations in the waterlevel impact the bathymetric maps since water depth dependson whether lake sonar measurements were collected in Juneor August.

A total of 28 lakes were surveyed. As shown in Table 2,eight lakes were surveyed near Barrow in year 1 (100 serieslakes), nine near Atqasuk in year 2 (200 series lakes) and 11near Reindeer Camp in year 3 (300 series). Lake area,length and perimeter were determined from the shorelinepolygon digitised from the CIR aerial photograph or DEM(see Table 2). The maximum measured (sonar) depth isindicated and differs slightly from the maximum rasterdepth based on the interpolation algorithm. The mean rasterdepth is also reported and represents the average depth ofthe entire lake. Owing to extremely windy conditions, lakesL104, L222 and L304 were incompletely surveyed. A sonartransect was conducted across the centre of the lake, so themaximum observed depth is likely representative. However,the transects did not cover the entire lake at a sufficientlythorough sampling density to justify interpolation, so noraster-based summary data are reported for these cases andno bathymetric maps were produced.

RESULTS AND DISCUSSION

Lakes on the OCP

The lakes surveyed on the OCP near Barrow (100 series)have maximum depths ranging from 1.5 to 2.6m, whilemean depth varies from 1.2 to 1.8m across the entire basin(Table 2). These lakes have a fairly uniform basin mor-phometry. As shown in Figure 2a and b and associatedtopographic profiles, there are no well-developed shelves.Instead, water depth increases fairly rapidly and uni-formly within 10–20m of the shore, then gradually slopes

Permafrost and Periglac. Process., (2012)

Figure 2 (a–f) Colour-infrared aerial images, bathymetric maps and topographic profiles from typical lakes in the three study areas, with dotted line showing sonartrace in (a).Water depth colour palette on the bathymetric maps is identical, but note the difference in depth scale. Vertical scale on the topographic profiles is identical.

K. M. Hinkel et al.

downward toward the maximum depth near the basincentre. Since winter ice thickness is 1.5–2.0 m in this area,the depths are generally sufficient to support liquid waterbelow winter ice and a talik likely exists beneath thecentral basin of these lakes.

Copyright © 2012 John Wiley & Sons, Ltd.

The morphometry of these lakes suggests fairly uniformbed subsidence with permafrost degradation. Owing to thehigh ice content, there is relatively little clastic or organicmaterial released when the ground thaws or the low(0.3m) banks erode, and little material available for

Permafrost and Periglac. Process., (2012)

Figure 2 Continued

Arctic Lakes Bathymetry

redistribution by currents. The net input and release ofmaterials to the basin is less than the volumetric expansionrate from ground subsidence, so the basin gradually

Copyright © 2012 John Wiley & Sons, Ltd.

deepens. Because ice content decreases with depth asthe shallow ice wedges thin (Carson and Hussey, 1962;Lewellen, 1972), basin deepening slows with time.

Permafrost and Periglac. Process., (2012)

Table

2Characteristicsof

lakesforwhich

bathym

etricdata

werecollected

in2008

(L100series),2009

(L200series)and2010

(L300series).

Site

Latitu

deLongitude

Region

Elev.

(m)

Length

(m)

Perim

eter

(m)

Area

(ha)

Max.measured

depth(m

)Max.raster

depth(m

)Meanraster

depth(m

)

L100

71.2416

-156.7745

OCP

7.9

2280

6860

185.7

2.56

2.56

1.82

L103

71.1226

-156.3141

OCP

4.7

2610

6150

182.8

1.86

1.85

1.47

L104

71.1930

-156.5022

OCP

4.4

2490

6820

177.9

1.45

nana

L106

71.1745

-156.8968

OCP

7.0

3770

9420

370.0

2.49

2.45

1.32

L107

71.2739

-156.4973

OCP

1.6

1880

4650

127.0

2.07

2.02

1.60

L118

70.9562

-157.1767

OCP

12.1

420

1390

11.3

1.94

1.78

1.17

L119

70.9959

-157.0397

OCP

10.5

470

1950

8.4

2.21

2.12

1.24

L121

71.1163

-156.0021

OCP

5.6

720

2470

34.0

2.22

2.18

1.35

L200

70.4526

-156.9517

ICP

19.8

2080

7590

250.9

2.59

2.53

1.13

L201

70.3294

-156.8378

ICP

29.0

1540

4900

147.9

2.56

2.53

1.02

L202

70.2879

-156.9849

ICP

29.5

1820

5230

153.0

3.35

3.26

1.18

L204

70.3727

-156.9613

ICP

26.5

1180

4140

42.9

2.53

2.51

0.83

L205

70.3773

-156.9267

ICP

28.3

2200

5580

157.9

2.33

2.29

0.99

L206

70.4180

-156.9833

ICP

25.8

1770

6290

189.3

3.08

3.06

1.22

L207

70.3244

-156.5886

ICP

28.5

3150

8680

358.9

3.70

3.33

0.78

L221

70.4363

-157.4047

ICP

22.0

2730

8090

287.8

5.22

5.12

1.13

L222

70.4198

-157.4205

ICP

22.4

2690

7110

244.8

5.01

nana

L300

69.9609

-156.5464

Yedom

a51.0

1410

3720

65.1

6.48

6.45

2.28

L302

70.0487

-156.8001

ICP

40.0

1040

3500

46.9

2.11

2.08

0.98

L303

70.0899

-156.7158

ICP

39.4

940

3050

46.9

1.81

1.77

0.86

L304

70.0599

-156.8302

ICP

40.7

2300

7530

211.1

1.32

nana

L305

70.1088

-156.7322

ICP

38.1

1020

3270

47.8

1.85

1.81

1.03

L306

69.9964

-156.5294

ICP

38.5

960

2650

46.0

2.82

2.80

1.53

L308

69.9864

-156.4252

Yedom

a54.0

1250

4080

80.2

8.05

7.54

2.52

L309

70.0244

-156.5671

ICP

37.6

1130

3210

61.3

3.85

3.52

1.75

L310

70.0144

-156.7026

ICP

43.8

1870

6390

215.5

2.30

2.28

1.49

L311

69.9955

-156.6895

Yedom

a56.5

1350

3770

81.1

9.52

9.16

3.11

L312

69.9533

-156.6389

Yedom

a57.5

1120

3860

81.7

8.18

8.00

2.29

OCP=Outer

Coastal

Plain;ICP=InnerCoastal

Plain;na

=no

spatialinterpolationdueto

incompletesurvey.

K. M. Hinkel et al.

Copyright © 2012 John Wiley & Sons, Ltd. Permafrost and Periglac. Process., (2012)

Arctic Lakes Bathymetry

Lakes on the ICP

Further inland, lakes on the ICP tend to have low banks(0.3m) and prominent littoral shelves composed of massiveor layered ice-bonded sand at shallow (0.4–0.7m) depths.Two examples are shown in Figure 2c and d. Lakes nearAtqasuk (200 series) ranged in maximum depth from 2.3 to5.2m, and mean depth averaged from 0.8 to 1.2m acrossthe entire lake basin. Although lakes here have greatermaximum depth, the mean lake depth is notably lower thanlakes on the OCP. This is due to the broad, shallow shelves,which often occupy a significant portion of the lake area. Inthis somewhat hilly terrain, bluff erosion yields copiousamounts of sediment for redistribution in the basin, and hasthe effect of reducing the volumetric capacity of the basinand decreasing the average lake depth. Although the shelveslack a talik (Burn, 2005), it is likely that fairly deep taliksunderlie the basin centres. Typically, sandy deposits are lesssusceptible to ice-enrichment processes and, although icewedges are present, there is little excess ground ice in theform of lenses, nets or veins; thawing sand does not experi-ence subsidence to the same degree as the ice-rich siltdeposits. Thus, to achieve the observed maximum depthsin the lakes, the talik must either extend deep beneath thelake bed or penetrate an ice-rich unit beneath the sandyaeolian deposit.

Lakes near the ACP-Arctic Foothills Transition Zone

Lakes near the ACP-Arctic Foothills transition zone show abimodal pattern of depth. Those lakes on the ACP side ofthe boundary have a maximum measured depth of 1.3 to3.9m, with mean lake depth ranging from 0.9 to 1.8 m.

Figure 3 Lakes developed on yedoma in the Arctic Coastal Plain-Arctic Foothillinto the upland feature. Blue is deep water, yellow is shallow. This figure

Copyright © 2012 John Wiley & Sons, Ltd.

Thus, they are similar to lakes on the ICP near Atqasuk interms of their maximum and mean depths, bank heightand basin morphometry.

Lakes developed on Yedoma

By contrast, lakes L300, L308, L311 and L312 are developedon yedoma in the silt belt and demonstrate markedly differentcharacteristics (Figure 3). The maximum measured depthsrecorded in these lakes range from 6.5 to 9.5m, andmean lakedepth from 2.3 to 3.1m. As shown in Figure 2e and f, most ofthe lake basin is morphologically similar to those found onthe OCP near Barrow, with low banks, a rapid increase inwater depth near the shore and gradual deepening towardthe central basin. However, deep pools are found in all fourlakes, and all pools are located near the margin of the lakewhere it is actively encroaching on a hill or ridge (Figure 3).Confined to a fairly small area some 15–50m from the shore,the deep pools are discontinuous and tend to parallel theshoreline.

Lake expansion and encroachment into the uplands areevidenced by active slumping on the bluff face, debris-flowfeatures, lack of vegetation on the bluff, the presence of fi-brous organic mats and shrubs in nearby water, and disco-louration of water at the lake margin. It is also clear fromFigure 3 that these lakes are remnants of larger lakes thatpartially drained to form an alas, and that the lakes are cur-rently expanding into the flanks of the upland feature. Giventheir proximity to the bluff, it is somewhat surprising thatthe pools have not been infilled with sediment.

At L300, we observed an ice- and organic-rich silt unitunderlying aeolian sand deposits at the base of the 20-mhigh northern bluff. The contact was about 1m above the

s transition zone showing the association of deep pools with encroachmentis available in colour online at wileyonlinelibrary.com/journal/ppp

Permafrost and Periglac. Process., (2012)

K. M. Hinkel et al.

current lake level and covered by debris except for a section2–3m wide. The upper unit is well-sorted silty sand andvery fine sand with cross-bedded structures, which gradesupward into massive silt. We interpret this to be a smalldune that was buried by loess. Below the sharp contact isorganic rich, grey sandy silt that is extremely ice-rich andexhibits ataxitic cryostructures. This layer extends to anunknown depth beneath the lake.Carter (1988) describes an exposure near the Reindeer

Camp study area that is 51m high. Thermoluminescencedates from the fluvial-aeolian sand interface near the baseof the exposure indicate that aeolian deposition beganaround 85 ka. Further upsection, loess accumulationbegan to dominate aeolian sand accumulation after around35 ka. At nearby exposures, broad ice wedges describedas ‘enormous’ exceeded depths of 15m, and ice contentin aeolian silt was typically more than 70 per cent(Lawson, 1988).At L300, it appears that loess deposits are vertically

continuous with older deposits that extend below the lakesurface and form a thick yedoma sequence. Pond andlake formation in the early Holocene likely began in localterrain depressions (‘depression’ thermokarst lakes ofJorgenson and Shur, 2007). Lake expansion and deepen-ing created a talik and thawed the underlying ice wedges.These processes halted when the lake was breached andpartially drained to create the large lake basins (alases)apparent in Figure 3. Talik penetration continued beneaththe remnant lake, and thaw subsidence yielded lake basinsthat are currently 2–3m deep. Eventual lake encroachmentinto the margin of the uplands removed the insulatingoverburden by bluff erosion, and lateral talik expansionsubsequently penetrated into the susceptible depositsbeneath the upland feature (Kokelj et al., 2005). As theice-rich sediments thawed, the lake bed surface was differen-tially lowered to a degree determined by the local icecontent. Near the bluffs, relatively recent thaw subsidencecurrently outpaces trough infilling by sedimentation. Thehigh-frequency variability apparent in the sonar transectsacross this part of the lake basin suggests melting of icewedges, but the resolution is inadequate to map thesefeatures from the lake surface. All four lakes developed onyedoma demonstrate the same pattern (Figure 3) and, in allcases, deep pools are found only where the lake is activelyeroding the upland feature.

CONCLUSIONS

Based on the variety of lakes that were surveyed in the threestudy areas, we conclude that the morphometry of lakesdiffers between the study areas and, as a first approximation,depends largely on the characteristics of the substratebeneath the lake.On the OCP, lakes deepen quickly near the shore and

gradually slope toward the basin centre. They are of fairlyuniform depth, with poorly developed littoral shelves. The

Copyright © 2012 John Wiley & Sons, Ltd.

maximum depth is around 2.5m, while the mean depth overthe entire basin is ~ 1.5m. Here, the uniform subsidence ofice-rich silt combined with the paucity of clastic sedimentsreleased after ground thaw yield uniform basin morphome-try. Since the ice content of the permafrost declines withdepth as shallow ice wedges taper out, the susceptibility tothaw subsidence decreases with time. There is relativelylittle sediment generated by lake margin erosion in thislow-relief terrain.

By contrast, lakes developed in the ice-poor aeolian sanddeposits of the ICP tend to have greater maximum depths(3–5m) near the lake centre, but exhibit lower mean basindepths of ~ 1.0m. Prominent littoral shelves develop fromredistribution and deposition of sand derived from erodingbluffs in this region of higher relief, yielding lakes of lowaverage depth. The greater maximum depths suggest thatthere has been deep penetration of the talik into the frozensediments beneath the deeper central parts of the lake. Inthese lakes, the gradual subsidence and increase in volumet-ric capacity associated with basin-centre talik developmentare offset by the input of clastic sediments from the lakemargin bluffs during lake expansion. These patterns holdtrue for all lakes studied on the ICP including those nearReindeer Camp.

Lakes developed on yedoma in the ACP-ArcticFoothills transition zone show a combination of character-istics. Most of the lakes are of fairly uniform depth, withrapid deepening near the shore and gradual deepeningtoward the basin centre. In this sense, they are morpho-logically similar to basins on the OCP that have alsodeveloped in ice-rich silt. That part of the lake encroach-ing on the flanks of loess hills and ridges has pools thatcan be up to 9m deep, and these features appear to resultfrom differential subsidence of the lake bottom above icewedges near the base of the yedoma. Presumably, thesedepressions are eventually infilled with clastic sedimentsderived from the erosion of lake bluffs. Given the overallcharacteristics of the lakes in this area, it is likely that theinitial lake basins developed directly on ice-rich depositsand were quite deep. Where thick loess deposition occurredin the upper part of the yedoma sequence, the underlyingice-rich deposits were insulated and protected until theoverburden was removed by more recent lateral lake expan-sion. These basins have a maximum average lake depth ofaround 2.5m, despite the significant input of clastic sedi-ments from the tall, actively eroding bluffs. The generalpattern of increasing maximum lake depth from the OCPto the ICP and ACP-Arctic Foothills transition zone isconsistent with the bathymetric estimates of drainedthermokarst lake basins identified by Wang et al. (2012).

The bathymetric data presented here can be used todevelop and calibrate models that use high-resolutionpassive or active sensing systems to estimate water depth.Remotely sensed data, supplemented with measured andmodelled maximum lake ice thickness, have been appliedin this area of the ACP to estimate general lake bathymetry(Mellor, 1994; Morris et al., 1995; Jeffries et al., 1996;

Permafrost and Periglac. Process., (2012)

Arctic Lakes Bathymetry

Kozlenko and Jeffries, 2000), but it is most effective forwater depths of less than 2m. Further, the coarse spatialresolution limits the ability to detect deep pools and steepgradients. Although the field results generally agree withthose derived from some of the more robust dynamicalmodelling efforts (e.g. Plug and West, 2009), they do high-light the need to incorporate the effects of sediment erosionand transport (Pelletier, 2005). The results suggest that thecollection of baseline data from a geographically representa-tive sample of lakes is essential to a more complete under-standing of the fundamental characteristics and processesof thermokarst lakes.

Copyright © 2012 John Wiley & Sons, Ltd.

ACKNOWLEDGEMENTS

This work was supported by grants from the National ScienceFoundation (KMH: 1107607, 0094769, 0713813; JDL:0713822; YS: 0013903; WRE: 0548846). Any opinions,findings, conclusions or recommendations expressed in thismaterial are those of the authors and do not necessarily repre-sent the views of the National Science Foundation. We aregrateful for the logistical support of Barrow Arctic ScienceConsortium, CH2MHill Polar Services and the UkpeagvikInupiat Corporation, and for the helpful remarks from thereviewers and Editor.

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