Lake-Level Fluctuations

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1		Edward Elling Berg	--	4851	2022-04-25 07:22:51	\|
2	update layout and references	Rita Xu	-5 word(s)	4846	2022-04-25 08:48:31	\|

This entry is adapted from the peer-reviewed paper 10.3390/quat5020023

Recent decades of warmer climate have brought drying wetlands and falling lake levels to southern Alaska. These recent changes can be placed into a longer-term context of postglacial lake-level fluctuations that include low stands that were as much as 7 m lower than present at eight lakes on the Kenai Lowland. Closed-basin lakes on the Kenai Lowland are typically ringed with old shorelines, usually as wave-cut scarps, cut several meters above modern lake levels; the scarps formed during deglaciation at 25–19 ka in a kettle moraine topography on the western Kenai Lowland. These high-water stands were followed by millennia of low stands, when closed-basin lake levels were drawn down by 5–10 m or more. Peat cores from satellite fens near or adjoining the eight closed-basin lakes show that a regional lake level rise was underway by at least 13.4 ka. At Jigsaw Lake, a detailed study of 23 pairs of overlapping sediment cores, seismic profiling, macrofossil analysis, and 58 AMS radiocarbon dates reveal rapidly rising water levels at 9–8 ka that caused large slabs of peat to slough off and sink to the lake bottom.

lake level change hydroclimate peat ice-shoved rampart

1. Introduction

Recent climate warming in the Arctic and Subarctic is well documented and is of international concern as a harbinger of the Earth’s future climate ^[1] (IPCC 2014). Climate records from southern Alaska, for example, closely track the climate variability of the broader North Pacific region, as reflected in indices such as the North Pacific Index (NP ^[2]), the Pacific Decadal Oscillation (PDO ^[3]) and the Aleutian Low Pressure Index (ALPI ^[4]). Hydroclimate, is dictated by fluctuations in these key modes of North Pacific climate variability and is of increasing interest for urban water supplies, agriculture, and forestry.

On longer timescales, hydroclimate variability also bears on the global carbon budget through its impact on carbon storage in peatlands. Moisture flowing from the North Pacific has created a vast storehouse of carbon in boreal and Subarctic peatlands since the last glacial maximum (LGM) at 26.5–16 ka ^[5]. Recent studies have suggested, however, that postglacial moisture flow into southern Alaska did not accelerate until several thousand years after the melting of the Laurentide and Cordilleran ice sheets was well underway. Cold water persisted in the Gulf of Alaska (GOA) until the Bølling warming at 14.7 ka ^[6], reducing moisture flow into southern Alaska ^[7]. Warmer sea surface temperatures (SSTs) in the GOA and rising sea level after Meltwater Pulse 1A (14.7–13.5 ka) brought moisture into Alaska from both the GOA and the flooding Bering Sea, initiating peatland recruitment ^[8], major vegetation change ^[9], and rising lake levels.

The entry presents a paleohydrological record for the central Kenai Peninsula since the end of the LGM, underscoring the impact of century- to millennial-scale modes of North Pacific variability. Water levels in closed-basin lakes (lakes with no outlets) closely reflect effective moisture (precipitation minus evapotranspiration), resulting from changes in local precipitation and evaporation from lake and soil surfaces, as well as evapotranspiration from catchment vegetation ^[10]^[11]^[12]. Here, researchers develop a record of lake-level fluctuations as a proxy for effective moisture using stratigraphic and geomorphic records, centered on Jigsaw Lake in the kettle moraine belt northeast of Sterling (Figure 1).

Figure 1. Map of the Kenai Lowland on the Kenai Peninsula. 1. Island Lake, 2. Pollard Lake, 3. Bay Lake, 4. Shadura Lake, 5. Elephant (Spirit) Lake, 6. Middle Finger Lake, 7. Horsetrail Fen ^[13], 8. Lake 93T, 9. Doroshin Lake, 10. Lake 79T, 11. Lake 78T, 12. Savka Lake, 13. Lake 88T, 14. Big Mink Lake, 15. Porcupine Lake, 16. Duck Lake, 17. W of Swan Lake, 18. Leaf Lake, 19. Portage Lake ^[14], 20. Mallard Lake, 21. Teal Lake, 22. S of Jigsaw Lake, 23. Arrow Lake ^[14], 24. Taiga Lake, 25. Coyote Lake, 26. Embryo Lake, 27. Trumpeter Lake, 28. S of Antler Lake, 29. Diamond Lake, 30. Moon Lake, 31. Aspen Lake, 32. Hidden Lake ^[15]^[16], 33. Barabara Lake, and 34. Kelly Lake ^[16]. Studies at labeled sites include Swanson Fen ^[17], Discovery Pond ^[9], Rainbow Lake ^[18], Paradox Lake ^[14], and Sunken Island Lake ^[19]^[20].

The Study Area

The Kenai Lowland (elevation 30–90 m asl) lies west of the Kenai Mountains (elevation 1000–1500 m asl), which form the spine of the Kenai Peninsula, separating Cook Inlet to the west from Prince William Sound (Figure 1). The mountains, comprised of Mesozoic bedrock, have experienced multiple Pleistocene and earlier glaciations. The Kenai Lowland is mantled by thick deposits left by glaciers from both the Kenai Mountains to the east and the Alaska Range to the west of Cook Inlet ^[21]. Glacial deposits range from highly permeable, ice-marginal sands to sandy tills, to relatively impermeable clay-rich cobble tills. A layer of fine sand- to silt-rich loess intercalated with numerous tephras mantle the uplands, ranging in thicknesses from centimeters to meters. The glaciated Kenai Lowland is underlain by flat-lying, Tertiary-age terrestrial sandstones, siltstones, shale, and coal, which are not exposed in the study area (Figure 1).

The study area is located on a rolling kettle moraine plateau (with elevations of 55 to 85 m asl) of Wisconsin age deposited during the Moosehorn stade, which reached its maximum between 23 and 19 ka ^[21]. This stade was by far the most extensive of the several LGM advances on the Kenai Peninsula ^[22]^[23]. At this climax of the LGM, Moosehorn ice advanced eastward from the Alaska Range, joined ice coming down Cook Inlet from the north, and extended onto the Kenai Lowland as far east as Sterling (Figure 1). A much smaller coeval advance extended westward out of the Kenai Mountains 30 km, terminating at the east side of Sterling. A 50 km-long glacial lake trending SW–NE covered today’s Moose River flats and separated the two lobes of Moosehorn ice. The eastern advance appears to have retreated without interruption, leaving few kettles on its ground moraine. The greatly extended western advance, however, left two terminal moraines, and then appears to have stagnated, leaving a jumbled, chaotic topography of eskers, kames, and kettles. This classic kettle moraine landscape shows many signs of high-water erosional activity such as wave-washed terraces, wave-cut scarps, underfit streams and steep-sided abandoned drainage channels ^[21]. The moraine and its surface and groundwater are drained internally by the westward-flowing Swanson River into Cook Inlet, on the western side by the south-flowing North Fork Beaver Creek (No Name Creek, local name), which drains into the Kenai River east of Kenai, and on the southeastern side by the Moose River which drains into the Kenai River at Sterling. The eight kettle lakes are distributed across the western Moosehorn moraine, aligned in a generally east–west direction. The Moosehorn glacial stade was followed immediately by the Killey stade (19–18 ka) with the recurrence of smaller glacial advances.

No evidence of permafrost effects, either ancient or modern, was observed in the study area, although actively degrading relict permafrost occurs in glacial lakebed wetlands of the Moose River Flats 10–15 kms east of the study area ^[24]^[25].

Jigsaw Lake, composed of four separate lobes, prompting the name “Jigsaw”, is located on a low drainage divide between the Swanson River to the north and the Moose River to the south (Figure 2). It has a relatively small catchment area (400 ha) compared to the lake itself (48 ha). Its basin is topographically closed, and ringed by a prominent wave-cut scarp at 85 m asl, 2–5 m high, and situated 2.4 m above the modern lake level.

Figure 2. Jigsaw Lake shown in a bare-earth LiDAR image, with peat coring locations in satellite fens and the surrounding Moosehorn moraines (yellow dots); lake sediment core locations are numbered dots within the lake basin. The yellow 85 m asl contour follows the wave-cut scarp of the ancestral lake basin.

The fact that the lake is situated at the top of its drainage divide potentially makes its water level more sensitive to variation in climatic effective moisture than lakes further down the watershed that would be buffered with groundwater flow from higher lakes ^[26].

The Kenai Mountains cast a strong rainshadow over the Kenai Lowland, capturing much of the precipitation coming off the Gulf of Alaska through Prince William Sound. On the windward (eastern) side of the mountains, there is heavy precipitation at Whittier (5050 mm/yr, mean annual temperature 4.4 °C) and Seward (1690 mm/yr, 4.3 °C). On the leeward side at Sterling (445 mm/yr, 1.4 °C) near the study area, and at Kenai (480 mm/yr, 1.3 °C), the climate is drier and more continental, and more similar to that of Interior Alaska ^[27]^[28].

At a regional scale, the climate of southern Alaska is driven by shifting patterns of the wintertime Aleutian Low pressure system. When the Aleutian Low pressure system is strong, it is centered over the Gulf of Alaska, and cyclonic storms frequently carry winter precipitation into south-central Alaska. With weaker Aleutian Lows, the low is elongated or splits between a west node near the western end of the Aleutian Islands and an east node in southern Prince William Sound; the west node routes winter storms toward the Alaskan Interior and the east node routes storms toward the Yukon and Southeast Alaska ^[4]^[29]^[30]^[31]^[32]. These shifting low pressure areas are partially tied to North Pacific sea surface temperatures described by the irregular 20 to 40 yr cycles of the Pacific Decadal Oscillation ^[3].

2. Jigsaw Lake Transported Peats

The multiple, discontinuous peat layers in Jigsaw Lake appear to be transported slabs. That these peats are not in situ is clearly indicated by slumped lake sediments in the seismic profiles and core stratigraphy that includes heavily fragmented peat textures, inverted dates, and sharp juxtaposition of unhumified plants and well-humified peats. The peats initially began accumulating prior to 10.4 ka on the lake shore apron that became exposed when the lake level had fallen almost 9 m below the 85 m asl wave-cut scarp. The peats were undercut by a rising lake level beginning approximately 9.2 ka (as shown by sloughed peat in cores #51 and DK 2001) and continued to slough off until 8 ka. A second period of rising lake level occurred approximately 4.5 ka and displaced more peat.

The satellite fens around Jigsaw Lake began accumulating peat during the initial lake level rise, i.e., 11 to 9 ka, even though they were perched several meters higher than the lake level and hence were not hydrologically connected to the lake at that time.

3. Jigsaw Lake Satellite-Fen Peat Coring

Fen D-D1 has the deepest and most complete record of lake level rise at Jigsaw Lake. Researchers interpret the lowest peat as a classic wet-to-dry sequence, starting with D. exannulatus growing at the lake water edge. Somewhat drier Sphagnum Sect. Cuspidata appears next, and is succeeded by dry S. fuscum and Calliergon stramineum, a generalist moss that can be found in both wet and dry habitats. This initial wet-to-dry sequence plausibly represents autogenic succession ^[33], as the fen surface has by now (~9.8 ka) grown an additional 1 m above its original water table; alternatively, the wet-to-dry shift may correlate with an interval of low stand at Dolly Varden Lake (Site B, discussed below). Wet conditions returned to satellite-fen D-D1 at the 400 cm depth (8 ka), with Sphagnum Section Cuspidata and the appearance of cladoceran carapaces; conditions remain wet through the remainder of the core. This upper sequence suggests, first, that the lake level has risen substantially and has nearly caught up with the Sphagnum surface at 8 ka, and second, that the rising peat surface generally tracks the rising lake level after 8 ka.

The full set of Jigsaw Lake fen peats is displayed in an elevation transect around the lake. The earliest vegetation in all the fens was composed of graminoids in shallow water, which then transitioned into dry Sphagnum fuscum. This transition is consistent with both autogenic succession and a mid-Holocene dry period (discussed below). Indications of wetter conditions subsequently appear in all the fens, albeit idiosyncratically, at times ranging from 4.8 ka (Cove Fen) to ~3.3 ka (Fens E, P & C); Fen H (the highest) shifts to somewhat wetter conditions at 7.2 ka but still retains S. fuscum thereafter, mixed with graminoids. Researchers interpret these time-transgressive dry-to-wet transitions as responses to a wetter late-Holocene climate, strongly modified by local site geomorphic and vegetation conditions.

The initial onset of peat accumulation was likely to have been strongly conditioned by local factors. The peat recorders, so to speak, are turned on in these fens over a wide time span, from 11.3 to 7.2 ka; regionally the span is even greater, from 14.2 ka (Swanson Fen) to 8.2 ka (Donkey Lake) (Figure 3). There could be at least three explanations for these differences; first, foundered ice blocks of different size and burial depths could melt out at quite different times in this intensely kettled topography ^[34]. In north-central Wisconsin buried Wisconsin-age ice persisted at least 5000 years after retreat of the active ice ^[35]. Second, peat accumulation requires an effective moisture high enough that the rate of vegetative accumulation exceeds the rate of decomposition. If the moisture regime fluctuates sufficiently, accumulated peat, unlike lake sediments, can be lost to decomposition or wildfire. Third, topographic relief on the substrate can allow peat to accumulate in low spots long before higher areas, so that multiple cores in the same peatland exhibit a several-thousand-year range of basal dates ^[36]. Differentiating among these factors is difficult, but not necessary for this entry; once the peat recorders are turned on, their hydrological records are independent of the starting date. Additionally, once turned on, the great water-holding capacity of (especially Sphagnum) peat tends to promote the accumulation of additional peat.

Figure 3. Summary chronology of central Kenai postglacial lake level variation. Lake level drops are measured below wave-cut scarps, which are typically several meters above modern lake levels. Lake level heights are estimated with respect to modern shorelines for Discovery Pond ^[9] and Swanson Fen ^[17], with the 12 ka date adjusted from 11.3 ka, following ^[9]. Discovery Pond is shown in red for clarity. Climate periods follow ^[37]^[38]. July insolation for 60° N is from ^[39]. Black dots show radiocarbon dates; arrows show fall or rise of lake level. The label “DV” shows peat segments of the Dolly Varden core which are separated by gyttja.

4. Reconstructed Water Level Changes

Researchers summarize water level changes in the regional lakes with a series of interpretive hydrographs showing the lake level chronologies as constrained by erosional features (wave-cut scarps and wave-washed terraces), terrestrial fen peats, and ice-shoved-ramparts. The hydrographs are superposed in a composite graphic with data from the literature (Figure 3). Seven distinct phases of lake level history emerge out of this synthesis: (1) An initial wave-cut scarp-forming period, (2) a falling lake level period culminating in the lowest stands, (3) initial rising lake levels 15 to 11 ka, (4) early Holocene rising lake levels (first rapidly, then more slowly), (5) a mid-Holocene period of relatively dry conditions, (6) a late-Holocene period of still higher lake levels, (7) culminating in late 20th century declines.

(1) The WCS-forming period (21 to 19 ka). Researchers lake chronologies begin with the erosion of high-level shorelines of the ancestral lakes that first formed in these basins during the last major deglaciation. The lakes originated as part of a large kettle field that formed as rising temperature caused the stagnation and downwasting of a glacier that had advanced onto the Kenai Peninsula from across Cook Inlet. Buried ice was likely still present in many kettles at ~19 ka when waves began cutting scarps along lakeshores. The steepness of the WCSs indicates considerable fetch for wind-driven waves, which could not effectively be generated until most of the remnant glacial ice had cleared from the lake basins.

Although researchers have not been able to directly date wave-cut scarps, they hypothesize that the scarp-forming period coincides with the final stages of Moosehorn (late MIS 2) deglaciation. The earliest dates for deglaciation on the Kenai Lowland (21–19 ka) are from cosmogenic exposure dating of glacially transported granitic boulders near the study area ^[40].

Modeling of southern Alaska climate at the terminal LGM (21 ka) suggests that summers were warmer, and that the AL was stronger in winter than it is today ^[41]^[42]. A strong AL would advect moist winter air into southern Alaska from the Gulf of Alaska (GOA) ^[4], providing snowfall for the Moosehorn ice sheets. Warm summers, however, would ultimately lead to melting these ice sheets.

(2) Falling lake level period (19 to 15 ka). Evidence of falling lake levels is represented by the shoulder aprons around closed-basin lakes. These aprons extend downward from the WCS shorelines, passing below modern lake levels to the depths of the adjoining satellite fens with terrestrial peat deposits. For example, at Jigsaw Lake, the falling period is represented by the slope that extends downward from 85 m asl scarp to 76 m asl. This lake level decline opened up 70% of the ancestral lake basin to vegetative colonization. Similar declines of 4 to 10 m occurred in all eight lakes examined and are one of the most striking features of these lake geomorphologies and their hydrological histories.

Researchers suspect that these declines reflect a millennial-scale epoch of aridity in southern Alaska, caused by a cold SSTs in the GOA during the deglacial period. Sediment cores taken in the GOA and along the British Columbia coast show that NE Pacific waters were 4–5 °C colder between 19 and 14.7 ka than the early Holocene ^[6]. These authors propose that SSTs fell in response to repeated deliveries of freshwater from Columbia River outburst floods, sourced from glacial Lake Missoula, as well as by meltwater from coastal glaciers. According to ^[7], the late-glacial midlatitude jet stream carried North Pacific moisture into southwestern North America and supported large pluvial lakes, such as Lake Bonneville in Utah, leaving the Pacific Northwest relatively depauperate in moisture. At 14.7 ka, however, modeling by these authors suggests an abrupt reorganization of this circulation, starting with major melting of the Cordilleran and Laurentide icesheets, warming of the GOA SSTs, and shifting of moisture flow northward into the Pacific Northwest and southern Alaska.

(3) Initial rising lake levels (15 to 11 ka). By 14+ ka, the fen “peat recorders” were turned on, and lake levels had already risen sufficiently for satellite fens to begin accumulating peat. For most of the fens, researchers interpret this stage as shallow standing water with graminoids, i.e., sedges and cottongrass. In all fens (except Dolly Varden, discussed below), once peat begins to accumulate, it continues to do so, commensurate with rising lake levels.

The oldest fen peats in the study area date to 14.2 ka (Swanson Fen, below), but the onset of peat accumulation varies over a 7 kyr interval (14.2 to 7.2 ka) at individual sites. The initial appearance of peat, or at least its preservation, seems to be controlled by site-specific factors. Even within a single lake basin, at Jigsaw Lake, the onset of peat accumulation in the nine satellite fens examined varies from 12.8 to 7.2 ka. Given that so many millennia can pass before effective peat accumulation begins at a site, the oldest satellite fens (Birch Lake 13.4 ka and Kayak Lake 13.0 ka) are likely distant minimum ages for rising lake levels.

M. Chipman (pers. comm., 2020) analyzed a 3.5 m sediment core that bottomed in gyttja at Rainbow Lake; this core provided a diatom-based water-depth reconstruction that shows shallow water (~4 m depth) in Rainbow Lake at 13.2 ka. When the diatom- estimate water depth was added to accumulated sediment thicknesses, the combined records clearly indicate that lake level rise was underway by 13.2 ka. At Discovery Pond (Figure 1 and Figure 3), a rising water table deposited a peaty mud with fen macrophytes at 13.4 ka on top of a non-organic silt ^[9]. Swanson Fen (240 m to the east of Discovery Pond) provides the oldest peat-based evidence of increasing effective moisture; in this basin, shallow-water pond sediments were deposited at 14.2 ka over a clay substrate and persisted until 14.0 ka, when brown moss peat began to accumulate ^[17]. At 13.7 ka, Chara is the first vegetation to appear at Horsetrail Fen, indicating the presence of shallow water; sedges and brown mosses followed shortly thereafter ^[8]^[13].

Two processes were co-occurring during the deglacial period that could have increased effective moisture and lake levels in southern Alaska. First, the Bølling Interstade, a warm interval from 14.7 to 14.1 ka, originally described from Europe and the North Atlantic ^[43], was coeval with and likely contributed to the warming of the GOA, described above. Second, Meltwater Pulse 1A (14.7 to 13.5 ka) was associated with rapidly rising sea level and flooding of the southwesternmost half of the Bering Platform ^[44], which enlarged the source area for moisture coming into Alaska.

Opposing the increased moisture of a warm GOA and the flooding of the Bering Platform, Northern Hemisphere summer insolation was increasing from its low at 20 ka to its maximum at 10 ka (Figure 3), with warmer summers tending to increase evaporative loss and reduce effective moisture. Regional climate modeling by ^[45], however, indicates that sea-level rise was by far the dominant factor affecting the hydroclimate. In their 11 ka simulation, the flooded southwestern Bering Platform provides moisture for westerly winds onto the Kenai Peninsula that make October the wettest month, just as it is today. In this modeling study, the Kenai Peninsula shows a strong winter Aleutian Low south of the Aleutian Chain at 11 ka, with easterly winds every month with lower temperatures and heat fluxes, and greater precipitation, all of which imply low evapotranspiration, high effective moisture, and rising lake levels at 11 ka, in spite of near-maximum summer insolation.

(4)Early Holocene rising lake levels (11 to 8 ka). Lake level rise in the study area appears to be well underway in the early Holocene. This rise is especially striking at Jigsaw Lake, where peat began to accumulate by 10.4 ka on the lakeshore apron that was exposed by the 8.8 m post-LGM lake level decline. As noted, between 9 and 8 ka, some of this peat sloughed off in slabs or detritus, as wave action on the rising lake undercut the sandy substrate.

(5) Mid-Holocene reduced effective moisture (8 to 4.8 ka). By 8 ka, peat accumulation slows in Jigsaw Lake Fen D-D, and various Sphagnum species are added to the graminoids (Figure 7); at the Cove Fen, there is a sharp transition from graminoids to dry Sphagnum fuscum at 7.9 ka. Jigsaw Lake Fens E-H all show S. fuscum periods in the interval 9.2 to 4.8 ka, which could be interpreted either as the result of autogenic succession or a drier climate; at the Cove Fen, however, it is unlikely that a hummock-former such as S. fuscum would be growing autogenically close to the lake surface (prior to 4.8 ka). A drier climate would be consistent with ^[46], who argue for a negative PDO phase and/or a more La Niña-like Northeast Pacific during the mid-Holocene, on the basis of reduced biologic productivity in the GOA, and reduced effective moisture in SE Alaska, SW Yukon, and Interior Alaska. A midge-based July temperature reconstruction from Rainbow Lake showed 4 to 7 ka to be the warmest period in the Holocene ^[18], reflecting the middle Holocene thermal maximum that is thought to have occurred in eastern Beringia ^[47]. In the dataset, Dolly Varden Lake site B1 provides clear evidence of lake level decline, where gyttja deposition ceased by 4.9 ka and wet Sphagnum (Section Sphagnum) moss established on the former lake bottom.

(6) Late-Holocene rising lake levels (4.8 ka to 20th century). Most of the Jigsaw Lake satellite fens suggest a second pulse of lake level rise in the late Holocene, in which the wet-to-dry trend is reversed, with dry Sphagnum fuscum being replaced by graminoids and various wet Sphagnum species. This dry-to-wet transition occurs at the Cove Fen (4.8 ka), Fen D-D1 (4. 1 ka), Fen C (3.3 ka), Fen P (2.5 ka) and Fen E (3.5 ka), but not at Fen H, which is mostly dry, nor at Fen L which is continuously wet. The dry-to-wet transition at Fen C (3.3. ka) occurred when the fen surface was ~4 m above the modeled Jigsaw Lake maximum water level. This suggests that Fen C was independently tracking the same climatic increase in effective moisture that was causing a rise in the Jigsaw Lake water level. The other late-Holocene dry-to-wet fens (Cove Fen, P and E) have their fen surfaces within ~1 m of the Fen D-D1 surface, which researchers interpret as a fairly close approximation of the lake surface because of its very wet vegetation assemblage. This late-Holocene rising lake level phase is also seen in sediment core #59-Upper with peat sloughing into the lake at 4.8 ka, and likewise in core #60 at 4.4 ka. At Kelly Lake (25 km south of Jigsaw Lake), a diatom record suggests that the lake level rose to near or above its modern level by 5 ka ^[16].

Episodic late-Holocene high stands (5 ka to 20th century). The late-Holocene high stands of Kenai Lowland lakes are clearly represented by well-developed ISRs, mostly placed 1–7 m above modern lake levels. The ISRs researchers examined typically record multiple shove events within a given rampart, showing that once an ISR is emplaced, it serves as a backstop for future shove events near the same level. The long age range of some ramparts indicates that the lake levels have not risen above the rampart since the first recorded shove event. Researchers interpret these events as episodic extremes because they have not found stratigraphic evidence (such as gyttja, textural, or plant species changes) at correlative dates in nearby peat cores. Researchers assume that the ramparts are formed during spring break-up with snowmelt high water and strong winds that drive ice pans onto the beach, bulldozing up shallow lake sediments. If high water persisted into the summer, the ramparts would be eroded by wave action before the next winter. The relatively young ISR dates (<5.2 ka, with many <2.4 ka) are consistent with large late-Holocene lake level rises which probably eroded or submerged older ISRs.

There is abundant evidence that the late Holocene was a climatically dramatic period in the NE Pacific and southern Alaska. The climate appears to have become more El Niño-like with a positive PDO pattern, and a strong eastward AL ^[46], which would bring more moisture from the Gulf of Alaska. The shift to a strong AL is evidenced by an ¹⁸O record from Sunken Island Lake that shows an abrupt ~2‰ increase in δ¹⁸O between 5.5 and 4.5 ka (isotopes measured on biological silica) ^[20]; similarly, an increase of 8‰ from 3 to 0.6 ka was observed at Horsetrail Fen (isotopes measured on total organic matter) ^[13].

Increased moisture at this time is suggested by a sharp decline in fire frequency from the early- and mid-Holocene mean of 12 to 8 fires/1000 yrs during the late Holocene, as determined from sedimentary charcoal at Paradox Lake (Figure 1) ^[14]. Increased snowfall associated with the strong AL likely supported the Neoglacial advances recorded in the Kenai Mountains at 3.6 ka, 600 A.D., and during the Little Ice Age, from 1300 to 1850 A.D. ^[48]. Strong ALs can provide the early spring winds necessary for large ISRs to form. The North Pacific (NP) index of measures the strength of the AL ^[2]; on the Kenai Peninsula, March-May E-W zonal winds correlate with the NP at r = 0.8 and N-S meridional winds at r = 0.3 ^[49]. The concentration of ISRs within the last 2.4 kyrs is striking, especially considering that only the most landward (oldest) ramparts were sampled. The presence of ramparts older than 2.4 ka at 9 of the 15 lakes suggests that the period of 0–2.4 ka was not accompanied by higher lake levels, which would have removed the older ramparts, but possibly had stronger spring wind events. This increased storminess is consistent with work along the outer coast of the GOA using traumatic resin ducts in tree-rings to infer increased winter storminess ^[50] and combined ice core studies from Denali and Mt. Logan that show an increase in the strength of the AL starting in the mid-18th century ^[51].

(7) 20th-century lake level drops. A general reduction in effective moisture (calculated as annual P-PET) on the order of 60% occurred in the central Kenai Peninsula after the 1968–1969 drought ^[12], with subsequent wetland drying and extensive invasion by black spruce and dwarf birch ^[52], as well as the occurrence of drought stress in trees and a massive spruce beetle outbreak ^[53]^[54]. Lake levels dropped by ~1 m or more in many lakes by the 1990s, especially in closed-basin lakes (E.E. Berg, pers. obs.). These moisture deficit effects are especially pronounced on the western Kenai Lowland because of the strong rainshadow created by the Kenai Mountains, but they may also be an expression of the general pattern of climate change now underway in southern Alaska ^[55]^[56].

5. Comparison with Interior Alaska

Two studies from Interior Alaska provide an interesting contrast to the study in southern coastal Alaska. Birch Lake (on the Richardson Highway, hereafter “Birch-R”) ^[57] and Harding Lake ^[58] are situated on the unglaciated Tanana River floodplain, north of the Alaska Range. These studies used lake sediment characteristics to estimate lake level elevations. Both lakes experienced substantial lake level rise at the end of the LGM at 16–15 ka, which is consistent with many studies ^[59] indicating that the Lateglacial was both warm and wet north of the Alaska Range. Surprisingly, both lakes showed substantial (>10–15 m) drawdowns sometime after 14 ka, contrary to the Kenai records, which indicate increasing effective moisture after 14.2 ka. The low stand at Birch-R occurred at ~13.7–12.0 ka; the lake rose to its modern overflow level by ~9.8 ka. Harding Lake showed similar low or fluctuating levels between 14 and 9.4 ka, when it rose to its modern level. Examples of other Interior lakes are reviewed in ^[47], all of which show higher/rising lake levels around 9 ka.

The LGM-Lateglacial-Holocene transition in Interior Alaska was strongly influenced by the retreating Laurentide icesheet, increasing summer insolation, and especially the flooding of the Bering Platform ^[44]. Researchers suggest, in contrast, that southern Alaska was most strongly affected by the Bolling warming of the GOA after 14.7 ka ^[6], which brought a flow of moisture into southern Alaska that kept lake levels generally high thereafter. Moisture flow into the Interior increased but also became much more variable when the Bering Sea became available as a moisture source ^[7]^[45] The Lateglacial drawdowns at Birch-R and Harding Lakes in the Interior are likely an expression of this variability.

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