Glacimarine sedimentary processes, facies and morphology of the south-southeast Alaska shelf and fjords

Marine Geology, 85 (1989) 359-390

359

Elsevier Science Publishers B.V., Amsterdam - - Printed in The Netherlands

GLACIMARINE SEDIMENTARY PROCESSES, FACIES AND MORPHOLOGY OF THE SOUTH-SOUTHEAST ALASKA SHELF AND FJORDS ROSS D. P O W E L L 1 and B R U C E F. M O L N I A 2 1Department of Geology, Northern Illinois University, DeKalb, IL 60115 (U.S.A.) 2U.S. Geological Survey, Office of International Geology, National Center, MS917, Reston, VA 22092 (U.S.A.) (Received September 1, 1987; revised and accepted May 20, 1988)

Abstract Powell, R.D. and Molnia, B.F., 1989. Glacimarine sedimentary processes, facies and morphology of the south-southeast Alaska shelf and fjords. In: R.D. Powell and A. Elverhoi (Editors), Modern Glacimarine Environments: Glacial and Marine Controls of Modern Lithofacies and Biofacies. Mar. Geol., 85: 359-390. High precipitation from Gulf of Alaska air masses can locally reach up to 800 cm a-1. This precipitation on tectonically active mountains creates cool-temperate glaciation with extremely active erosion and continuously renewed resources. High basal debris loads up to 1.5 m thick of pure debris and rapid glacial flow, which can be more t h a n 3000 m a-1, combine to produce large volumes of siliciclastic glacimarine sediment at some of the highest sediment accumulation rates on record. At tidewater fronts of valley glaciers, sediment accumulation rates can be over 13 m a-1 and deltas commonly grow at about 106 m 3 a -1. Major processes influencing glacimarine sedimentation are glacial transport and glacier-contact deposition, meltwater (subaerial and submarine) and runoff transport and deposition, iceberg rafting and gouging, sea-ice transport, wave action and storm reworking, tidal transport and deposition, alongshelf transport, sliding and slumping and gravity flows, eolian transport, and biogenic production and reworking. Processes are similar in both shelf and fjord settings; however, different intensities of some processes create different facies associations and geometries. The tectonoclimatic regime also controls morphology because bedrock structure is modified by glacial action. Major glacimarine depositional systems are all siliciclastic. They are subglacial, marginal morainal bank and submarine outwash, and proglacial/paraglacial-fluvial/deltaic, beach, tidal flat/estuary, glacial fjord, marine outwash fjord and continental shelf. Future research should include study of long cores with extensive dating and more seismic surveys to evaluate areal and temporal extent of glacial facies and glaciation; time-series oceanographic data, sidescan sonar surveys and submersible dives to evaluate modern processes; biogenic diversity and production to evaluate paleoecological, paleobiogeographic and biofacies analysis; and detailed comparisons of exposed older rock of the Yakataga Formation to evaluate how glacial style has evolved over 6.3 Ma.

Introduction

The glacimarine setting is influenced by each of Earth's sedimentological transport agents: ice, water and wind. Situations vary from where the glacial agent is dominant to 0025-3227/89]$03.50

where water is dominant with ice as a subordinate influence. Ice is subordinate today in the north-northeast Gulf of Alaska region on the continental shelf and in marine outwash fjords (Fig.l), because glaciers have retreated from the sea and glacially fed rivers contribute

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361 sis is on the modern, so t h a t products can be attributed to known processes; however, discussion of glacial action on the continental shelf necessitates occasional brief discussions of older Cenozoic deposits. Some of the major processes and controls that act in fjords and continental shelf of south-southeast Alaska are poorly known. Certain types of facies and their distributions through time are equally poorly understood. In fact, no continuous cores have penetrated the shelf to sample the forearc basin sediments and so we do not even know the maximum extent of Wisconsinan glaciation (compare Hamilton and Thorson (1983) with Molnia (1986)), which is important for ice-volume estimates and its effect on eustatic sea level (cf. Meier, 1984). Likewise, information for reliable lithofacies and biofacies modelling is scant, which reduces its utility for assisting interpretations of ancient sequences.

marine outwash gravel, sand and glacial rock flour. Both shelf and fjords were under a stronger glacial influence during the Wisconsinan but only glacial fjords, those with glacial tidewater fronts, are dominated by ice today. Water, in the form of seawater and streams is the major modifying agent once debris is released from ice. Wind influence is relatively minor in the region compared with the other two agents. The Gulf of Alaska region has a relatively mild climate compared with other glaciated areas today and the same relationship occurred during the Wisconsinan (CLIMAP, 1976). Consequently, it is an important end member in the spectrum of glacial regimes. Tectonism causing orogenesis and favorable continental orientation to ocean water masses and air masses, is the primary control over glaciation. This setting allows glaciers to extend to sea level at the relatively low latitude of about 57°N. Chile is under a similar tectonoclimatic regime and is the only region of the world at present where sea-level glaciation occurs at lower latitude. Along south-southeast Alaska, the cooltemperate climate supports many environments analogous to non-glacial settings, each superimposed with glacial domination or glacial influence. Much of the information available for describing the geological environment of the area was recently summarized by Hampton et al. (1987). The aim of this paper is to highlight what are considered to be some of the more important processes and products with respect to the glacimarine setting. The empha-

Controls o f sediment production First-order controls on glacimarine sediment production in south-southeast Alaska are tectonics, climate and sea-level changes. Fluctuations in sea level are most important to consider when interpreting stratigraphic successions such as when older Cenozoic deposits are discussed. Of the second-order controls on glacimarine sediment production in the region (Table 1), the glacial, marine, and fluvial are the most important. Fluvial controls are mainly glacial or climatic in origin and will be considered under those topics. Wind is a

TABLE 1 Major controls on glacimarine sedimentationin the Gulf of Alaska region Ice

Water

Wind

Biogenic activity

Glacier ice Icebergs Sea ice

Meltwaterstreams Precipitation and runoff Surface waves

Synoptic Local: Onshore-offshore,and katabatic

-

Seasonal water-column structure Baroclinic and barotropic currents Water-mass velocities

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climatic phenomenon and, similarly, will be considered under the first-order control of climate. Tectonic controls

The style of Cenozoic tectonism of the Gulf of Alaska region is one of plate convergence and terrane accretion (summaries in Howell, 1985; Jacob, 1987). This continuous orogenesis creates high-elevation accumulation areas for precipitation, and indeed, is a major forcing function for orographic precipitation. Continental configurations and latitudinal positioning are also tectonic factors and are important for forcing oceanic and atmospheric circulation, such that warmer air masses having travelled many hundreds of kilometres across the Gulf of Alaska are cooled and made to rise thousands of meters at the coastline. Tectonic uplift combined with locally fractured and metasomatized rocks create continuously renewable sources for glacial erosion and debris production. Mineralogy of marine clays shows they are mainly chlorite and kaolinite and that rock breakdown is primarily physical rather than chemical (Wright and Sharma, 1969; Molnia and Hein, 1982). Tectonics also influence general topography and sea-floor morphology. Active and dormant strike-slip faults and their splays separate tectonostratigraphic terranes of the southsoutheastern Alaskan margin (Jacob, 1987) and major valleys and fjords are aligned with these fault systems and other tectonic elements such as thrust faults. However, the troughs themselves must be erosional rather than simply tectonic in origin because they are currently still troughs. Vertical movements combined with isostatic rebound and high rates of sediment accumulation should lead to total infilling and shallowing of fjords. However, seismic reflection profiles show local basement depths to 600 m and a bottom sediment surface up to 500 m below sea level. This relief must be a consequence of relatively recent glacial erosion. In this type of tectonic regime, shelf succes-

sions can become quite thick, for example, in the Yakataga Formation glacimarine sediments are over 3 km thick (Plafker and Addicott, 1976; Armentrout, 1983; Eyles, 1987, 1988). These packages accumulate by tectonic deepening of a forearc basin and they can be preserved by seaward migration of the subduction complex and continued uplift of the orogenic arc (e.g., Dickinson and Seely, 1979), as well as by translation of terranes along strike-slip faults. Configurations like this have undoubtedly occurred in different places throughout earth history and glacimarine successions similar to those of the northnortheast Gulf of Alaska should predictably occur in ancient subduction and collision orogenic belts. Climatic controls

Climate is important in several ways because it controls glacial regime, volume of precipitation and runoff, organic processes and ecosystem distributions, oceanic temperatures, and circulation and wind-induced processes. Seasonality is an important consideration, as in other glacial regimes, for lithofacies types and distribution, and sediment accumulation rates. Weather systems generating precipitation and winds along the Gulf of Alaska coast are dominated by seasonal movements of the Aleutian low and North Pacific high-pressure cells (Sobey, 1980; Wilson and Overland, 1987). The moist maritime air from stagnant lows over the gulf create heavy precipitation that feeds the glaciers at high elevation during winter. Annual-mean precipitation along the coast is between 200 and 300 cm and precipitation reaches a maximum of over 800 cm a - ~ in the mountains of southeast Alaska (Wilson and Overland, 1987). At Yakutat, southern Alaska, the rainfall: snowfall ratio is about 1:2 (Sobey, 1980). Peak monthly mean precipitation occurs in winter (Reed and Elliot, 1979) but peak rainfall and runoff occurs in late autumn to early winter for coastal streams (NOAA, 1985), when most storms occur (Miller,

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1963). In addition, precipitation appears to increase with altitude (Murphy and Schamach, 1966), however, records are very poor and do not permit quantification of these systems. High rainfall combined with glacial meltwater produce high annual-mean freshwater discharges into the Gulf of Alaska, which are estimated at 2.4 × 104 m s- 1 (Royer, 1982). These discharges drive the coastal current of the Gulf of Alaska (Royer, 1979). Not only high precipitation favors glaciation of this coastal region, cloud cover is also important. Cloudiness is a factor in the balance of thermal energy flux to the surface of the gulf because the high cloud cover cuts annual sensible heat transfer and reduces glacial ablation. The monthly-mean cloud cover over the gulf ranges between 75 and 88%. Annual-mean sea-level air temperatures are between about 4 ° and 6°C with monthly means as low as 2°C in winter and as high as 14°C in summer (Wilson and Overland, 1987). Monthly-mean sea-level wind speeds are as low as 6 cm s- 1 in summer and as high as 11 m s 1 during winter, but the Aleutian low can propagate wind speeds greater than 25 m sduring winter (Wilson and Overland, 1987). These storms can stagnate over the gulf, being blocked by coastal mountains and a temperature-induced pressure gradient (Nummedal and Stephen, 1978). Ekman forcing moves near-surface water of the gulf toward the coast causing downwelling during most of the year, although during summer westerlies can locally move water offshore and occasionally may allow upwelling at the coast (Reed and Schumacher, 1987). Winds can also enhance mixing; greatest mixing occurs during winter because stronger winds break down stratification. Both the Ekman forcing and mixing are important for glacimarine sedimentation by influencing both the distribution of suspended particulate matter (SPM) and biological activity. Storms produce large waves with extreme heights of 40 m, but more commonly of about 13 m (Hampton et al., 1987). Such large waves can cause sediment erosion on the continental shelf as

well as propagate mass failure of sediment (Rappeport, 1980; Schwab and Lee, 1983). Katabatic winds flowing offshore influence coastal regions until the winds subside from convective warming by the ocean sometimes more than 25 km offshore. Fall or Taku winds can flow down major fluvial drainage systems (Alsek and Copper River valleys) and transport loess into the g u l f (e.g., Post, 1976). However, the loess is very diluted by other inputs, primarily from stream discharges, and so is relatively insignificant in the sediment record (Molnia, 1983). Glacial winds, funnelled into fjords from valley glaciers, influence the paths of icebergs and turbid overflow plumes (Cowan and Powell, 1986). The final consideration of climate is sea ice. It is included here because in the northnortheast Gulf of Alaska region, most sea ice does not form by freezing of sea water, but rather by freezing of large volumes of snow accumulated on the sea surface. This process requires quiet water and consequently, sea ice forms only at heads of fjords in winter. Some is fast ice that may detach with tidal fluctuations and break out as floes to form pan ice during wind storms resulting in minor ice rafting of debris. Sea ice can also be locally important on tidal flats (Bartsch-Winkler and Ovenshine, 1984).

Glacial controls Types of glacimarine lithofacies are very dependent on glacial budget, temperature gradient through a glacier, and basal temperature because these three factors control (1) type of glacial front, (2) rates of frontal advance and retreat, (3) debris budgets, that is, rates and volumes of debris supplied to the sea (erosion and glacial flow or transportation rates), and (4) melting rate, sediment loads and discharges from meltwater streams.

Type of glacial front At present, glaciers cover about 20~/o of the Gulf of Alaska coastal area (Royer, 1982) and more than 100 valley glaciers end in the sea as

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grounded tidewater cliffs under the cooltemperate to mild sub-polar climatic conditions. Development of floating glacial fronts (ice shelves) in temperate glaciers is inhibited because of tensile weakness due to water films between ice crystals. Rate of frontal movement Glaciers have been advancing and retreating through the fjord systems asynchronously with climatic change since retreat from the continental shelf after the Wisconsinan (Field, 1947, 1979; Goldthwait, 1966, 1986; Mann, 1986a). A change in mass balance is the first-order control on fluctuations of tidewater fronts (Mercer, 1961), however, other secondary factors are also seen as important for glacial stability in south-southeast Alaska. Sea-floor morphology, such as basins, banks and sills, and headland configurations, is significant because water depth has been shown to influence the stability of tidewater fronts (Brown et al., 1982), at least averaged over long time periods (Powell, 1984b). Subordinate factors of frontal stability are melting by seawater (cf. Weeks and Campbell, 1973), meltwater streams (Sikonia and Post, 1980), slope of the sea floor (cf. Thomas and Bentley, 1987; Denton and Hughes, 1981), and crevasse systems (Powell, 1983b, 1984b). Most glacial advances into the sea are commonly stow (e.g., Field, 1947, 1979; Goldthwait et al., 1963) because either a sediment threshold (morainal bank) needs to be constructed and/or sea level must fall in order to maintain a stable water depth at the tidewater front. However, some glaciers along the coast are known to surge (Meier and Post, 1969; Kamb et al., 1985) and some of them can move quickly in to the sea (e.g., Eliot, 1987). Past valley glaciers ending in fjords appear to have produced large volumes of sediment just as they do today. Some fjords, such as Glacier Bay, southeast Alaska have had sediment fills to more than 140 m above present sea level (Goldthwait, 1966, 1986). During intergtacials, these fjords act as sinks, trapping most sediment before it reaches the continental shelf.

The fills appear to have been almost totally eroded during subsequent glacial advance (Carlson et al., 1983; Molnia et al., 1984), and the sediment is presumed to have been transported to the shelf although appropriate budgets have not been documented. A pile of sediment at the base of the continental slope offshore from Cross Sound (Ludwig and Houtz, 1979) is also a potential site for the glacially reworked sediment from Glacier Bay. Historic retreat rates for tidewater fronts in south-southeast Alaska are well known (Field, 1947, 1979; Meier et al., 1980). After the initial climatic instability, retreat rates are controlled primarily by water depth (Post and LaChapelle, 1971; Brown et al., 1982). Tidewater fronts retreating at different rates have been shown to produce different glacimarine lithofacies associations (Powell, 1981, 1983b). These data lead to the conclusion that sedimentation appears to be episodic, being concentrated at depocenters that change locations through a glacial period. Migration of depocenters occurs because the glacial agent is unique compared with other agents in that the sediment source may move hundreds of kilometers in a virtual instant of geological time. Accumulation in fjords and on the continental shelf is asynchronous; it is high in fjords during interglacials, high on the outer shelLupper slope (and beyond?) during glacial maxima, and high on the shelf during glacial advances and retreats to and from each maximum.

Debris sources and budgets Most ablation of tidewater fronts occurs by calving, but rapid surface melting in the ablation zone produces, in some cases, significant volumes of supraglacial debris. That debris originates either by mass movement processes from high valley walls, or from basal debris when a tributary glacier joins a trunk glacier. A glacier's surface for a few kilometers back from the front is highly fractured by large transverse crevasses from extensional flow. Englacial and supraglacial debris accumulates

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at the base of these crevasses, which can be below sea level. Most supraglacial debris is dumped at the grounding line as icebergs calve. Debris in crevasses can slough off at the same time, but none are preserved as discrete flow units on the sea floor; sea-floor deposits are rock- or grain-fall accumulations that are incorporated into bergstone mud or diamicton facies, interstratified sand/silt and mud facies, or a morainal bank system (Powell, 1981, 1983a,b).

There is evidence t h a t these valley glaciers ending at tidewater contain moderate to large quantities of basal debris as do land-ending glaciers (cf. Lawson, 1979, 1981). Two sources can be examined for debris (Fig.2): icebergs calved from the base of active tidewater fronts which contain basal and englacial debris, and glaciers that have just retreated from the sea and which still retain vertical cliffs with exposed basal debris (Heiny and Powell, 1982). An example of such observations at Riggs and

Fig.2. Two examples of how basal debris can be studied at tidewater fronts, a. Exposures on fronts recently retreated from the sea. b. From icebergs calved from the base of the glacier.

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McBride glaciers, Glacier Bay show that on average those glaciers have debris zones that are concentrated in about the basal 14 m of the glacier. Of th at debris, the lower 6 in is stratified, about half of which consists of debris layers with average debris concentrations of 40% by volume in the ice. The rest of the stratified zone has dispersed debris with average debris concentrations of 3% by volume similar to the upper 8 m of the basal zone (cf. Lawson, 1979, 1981). Given these generalized estimates, glaciers ending as tidewater fronts may have about 1.5 m thickness of pure debris per square meter of grounding line. Subglacial processes of marine glaciers have not been measured in the area. Lodgement and perhaps local meltout tills in subglacial cavities are presumed to be deposited f r o m t h e base of glaciers while they end in the sea. However, no such deposits have been sampled from the sea floor during glacial r e t r e a t because sediment accumulation rates are too high. The best evidence for lodgement in the sea is a highly consolidated diamicton cored on an abandoned morainal bank at Reid Inlet, Glacier Bay (Powell, 1981, 1983b). Budgets of glacial debris have not been quantitatively determined. Melting-freezing processes are assumed to be very active at the glacier sole at high altitudes in order to produce the volumes of debris at termini. Dedundation rates over a 29 year period for a glacial drainage basin in Glacier Bay were estimated at 1.2 cm a -1 (McKenzie, 1986). Although not directly comparable with the estimated world average maximum rate of I m ka 1 (Schumm, 1977, p.34), the order of magnitude difference appears to be significant. Subaerial mechanical weathering and erosion are also very efficient and produce significant volumes of supraglacial debris t hat can be moved to englacial and subglacial positions during glacial transport. Furthermore, glacial flow velocities can be greater t han 3000 m a - 1 (Brown et al., 1982); thus a large debris input and rapid glacial flow combine to give high flux rates of debris to glacial termini in the sea.

Melting rates and streams Due to the maritime climate along the gulf coast with moderate air temperatures, glacial ablation rates are rapid. At calving margins ice toss to icebergs can reach about 109 m 3 a l (data in Brown et al., 1982). Calving is the dominant ablation process although melting is also important. Average melting rates for the region are unknown but, for example, glacial down-wasting in Glacier Bay has occurred at rates between 3.0 and 9.5 m a ~ since 1892, with an average of about 8.5 m a I (Goldthwait, 1986). Large discharges from meltwater streams attest to high melting rates but these discharges can be enlarged by high rainfall (e.g., Gustavson and Boothroyd, 1982). Runoff from deglaciated basins, valley walls, and from the glacier surface, can contribute significantly to meltwater streams. In the ablation zone rai nw at er can flow to the base of a glacier, a situation different from sub-polar and polar glaciers where surface streams are confined higher in the glacier. Sometimes subglacial streams respond virtually instantaneously to a precipitation peak (Cowan et al., 1988). Marine controls Continental shelf The Alaska Current dominates the hydrography of offshore Gulf of Alaska and flows at up to 30 cm s- ~ toward the n o r t h w e s t - w e s t at the edge of the south-southeast Alaskan continental shelf (Fig.l) (Reed and Schumacher, 1987). The continental shelf to the southeast is 20-50 km wide and is dominated by a horizontal salinity gradient created by coastal freshwater influx. This influx drives a geostrophic flow, the Alaska Coastal Current n o r t h w e s t - w e s t at surface velocities of between 10 and 20 cm s 1 (Reed and Schumacher, 1987). The coastal c u r r e n t flows along the inner shelf until it is forced out into the Alaskan Current by K a y a k Island protruding from the co~st. West of K ayak Island, where the shelf is 100 km wide, there is a large gyre and waters are dominated by the Copper River and surface velocities in

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the coastal current are more than 20 cm s(Reed and Schumacher, 1987). Lowest salinities of the coastal current in September and October correlate with maximum discharge from the land (Royer, 1979). Satellite images show the SPM from these discharges is transported in turbid surface plumes toward the northwest-west along the shelf in the coastal current and is eventually deposited in nearshore depressions and embayments (Reimnitz and Carlson, 1975; Molnia and Carlson, 1978). The terrestrial component of 80% of the SPM (Feely et al., 1979), is derived from glacial sources on land. Wave-climate studies confirm the net northwest-west alongshore sediment transport at an average rate of 2 × 106 m 3 a 1 (Nummedal and Stephen, 1978). Streams entering the gulf from Bering, Guyot and Malaspina glaciers are important sediment sources for the shelf. Turbid plumes from their discharges carry about 1 mg 1-1 suspended particulate matter (SPM) (Feely et al., 1979). Copper River is the largest contributor of SPM to the shelf, its turbid plume having a highest measured SPM concentration of 6 mg 1-1, and an average greater than 2 mg 1-1 (Feely et al., 1979). Some surface SPM becomes entrained in gyres and transported offshore such as west of K a y a k Island, between Icy Bay and Yakutat Bay, and east of Yakutat Bay (Carlson et al., 1977). The strong gyre west of Kayak Island transports 3000-24,000 t day 1 when it is active over summer (Feely et al., 1979). Average bottom-current velocities are lower than surface velocities; at 10 m above bottom the coastal current velocity is reduced to about 7 cm s 1 (data in Reed and Schumacher, 1987). Over mid-shelf regions bottom sediment can be resuspended to form nearbottom nepheloid layers, but the long-term trend favors net deposition (Feely et al., 1979). The Alaska Current or wind forcing are thought to cause prolonged events of highvelocity currents (up to 25 cm s-1 over several days) that are capable of resuspending significant amounts of bottom sediment and transporting it northwest-west along the shelf (Feely et al., 1979).

Near-bottom sediment is not transported offshore from the continental shelf between Yakutat and K a y a k Island; however, west of K a y a k Island dense near-shore water produced by winter cooling and wind-driven downwelling forms an offshore, gravitationally convective flow with bottom velocities up to 9 cm s- 1 and average sediment concentrations of 1.5 mg 1 1. Flow across the shelf is stopped by the Alaska Current at a depth of 180-200 m. Consequently, near-bottom sediment is transported offshore beyond the edge of the outer shelf and then dispersed (Feely et al., 1979). Mixed semi-diurnal tides propagate counterclockwise in the Gulf of Alaska and flow at 6-10 cm s 1 on the shelf (Sobey, 1980). Tidal currents produce local high-frequency fluctuations in near-bottom sediment concentrations, which are thought to be due to resuspension of bottom sediments followed by rapid redeposition with little net transport (Feely et al., 1979).

Fjords Two types of fjords have been defined in south-southeast Alaska based on their glacial character (Burrell, 1971). Glacial fjords have glacial tidewater fronts; turbid outwash fjords receive glacigenic sediment but glaciers end on land or are absent within the drainage basin. Glacial fjords have smaller ranges for salinity and density and lower water temperatures but higher dissolved oxygen content (Pickard, 1967). Distinct water masses are created in glacial fjords by mixing of two different freshwater sources with salt water: direct melting of glacial ice by seawater (dominant in winter) and glacial meltwater streams mixing convectively with fjord water (Greisman, 1979). Major stratification within a fjord water column is generated in summer and autumn by surface freshwater flowing seaward as a barotropic flow from streams either at a delta or a glacial tidewater front. Extreme heterogeneity occurs in October and November (Matthews and Quinlan, 1975). Katabatic winds flowing from glaciers at speeds up to 50 m s -1 (Matthews, 1971) enhance both upwelling at a

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glacier front and down-fiord wind stress of surface layers. Beyond the major d e p o c e n t e r s of t i d e w a t e r fronts and deltas, most sedimentation in the fjords is controlled by the hypopycnal flows. Particle settling from turbid plumes in upper w a t e r layers is d e p e n d e n t on plume dispersal, t u r b u l e n c e , sediment c o n c e n t r a t i o n s and sizes, and flocculation and pelletization (el. Syvitski et al., 1987). To b a l a n c e the h y p o p y c n a l outflow, an i n t e r m e d i a t e extrabasinal w a t e r layer e i t h e r from the c o n t i n e n t a l shelf or the n e x t basin down-fjord, flows in as a baroclinic current. Stratification of the w a t e r c o l u m n is decreased or e l i m i n a t e d in w i n t e r and m a x i m u m homogeneity commonly occurs during F e b r u a r y and M a r c h ( M a t t h e w s and Quinlan, 1975). The h o m o g e n e i t y is p r o d u c e d by a decrease in f r e s h w a t e r input, t h e r m o h a l i n e c o n v e c t i o n due to surface cooling a n d / o r seaice formation, and deep-water renewal. F u r t h e r m o r e , coastal winds can force extrabasinal w a t e r into fjord systems and if a fjord has no e n t r a n c e sill, large storm waves can be propagated. Deep w a t e r r e n e w a l o c c u r s when extrabasinal w a t e r flows over the e n t r a n c e sill by density c u r r e n t s t r i g g e r e d by spring tidal motion, p r o p a g a t i o n of i n t e r n a l waves, or baroclinic or b a r o t r o p i c c u r r e n t s g e n e r a t e d by wind (el. Syvitski et al., 1987). The m a x i m u m density of Gulf of Alaska w a t e r o c c u r s in w i n t e r at sills s h a l l o w e r t h a n 40 m, and then bottom-water r e n e w a l occurs. W h e n sills are deeper t h a n 100 m, r e n e w a l occurs d u r i n g w i n t e r t h r o u g h to s u m m e r ( M u e n c h and Heggie, 1978). Even in shallow-silled fjords, deepw a t e r renewal occurs at least once a year; therefore, no fjords in the region are s t a g n a n t (Reed and S c h u m a c h e r , 1987). Sedimentological c o n s e q u e n c e s of these r e n e w a l s are poorly u n d e r s t o o d but no a n o x i c sediment is p r o d u c e d and erosion and r e s u s p e n s i o n of b o t t o m sediments may o c c u r d u r i n g s u m m e r while some s e d i m e n t a t i o n may be p r o d u c e d locally d u r i n g w i n t e r (el. Syvitski et al., 1987). Tidal c u r r e n t s in fjords can be amplified beyond those on the shelf to obtain velocities

of over 5 m s ~ and spring tidal r a n g e s over 7 m. These c u r r e n t s may m a i n t a i n c o n t a c t with bottom sediment (above 20 cm s 1) and could produce sand-wave fields and lag surtaces ( H a m p t o n et al., 1987). Tides are also very i m p o r t a n t in c o n t r o l l i n g surface plume dispersal and as a c o n s e q u e n c e influence lithofacies close to t i d e w a t e r fronts as will be discussed in more detail u n d e r s e d i m e n t a r y products.

Sedimentary products G l a c i m a r i n e depositional systems along the n o r t h - n o r t h e a s t G u l f of Alaska can be broadly subdivided into t r a n s i t i o n a l and m a r i n e systerns (Fig.3). M a r i n e systems are t a k e n as those t h a t are c o n s t r u c t e d e n t i r e l y below mean lowest low-tide level in fjords and on the c o n t i n e n t a l shelf. The t r a n s i t i o n a l systems form along shorelines above m a r i n e systems but still within influence of the sea. W i t h i n the two systems, t h e r e are glacier marginal settings and proglacial and paraglacial settings. P a r a g l a c i a l was used by C h u r c h and R y d e r (1972) to define non-glacial processes t h a t are directly conditioned by glaciation and includes '~deposits [that] r a n g e from alluvial c o n e s . . . t h r o u g h alluvial fans to large alluvial plains and deltaic deposits in lakes or the sea" (Church and Ryder, 1972, p.3061). T h a t description is extended here to include m a r i n e systems t h a t receive glacigenic sedim e n t but are not in direct c o n t a c t with glacial ice (both glaciers and icebergs), t h o u g h sea ice may c r e a t e i m p o r t a n t processes. W r i g h t and S h a r m a (1969) used the term periglacial to describe these types of systems from the fjords and c o n t i n e n t a l shelf of the G u l f of Alaska. P a r a g l a c i a l is p r e f e r r e d because intense frost action is not c h a r a c t e r i s t i c of the region (cf. Pewe, 1983) and is n e v e r active in m a r i n e systems. C h u r c h and R y d e r (1972) include proglacial with p a r a g l a c i a l settings, however, the r e l a t i v e closeness of glacier m a r g i n a l e n v i r o n m e n t s with the proglacial setting creates distinctive processes in some instances and the d i f f e r e n t i a t i o n of proglacial and distal paraglacial is considered useful. Even with

369

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Fig.3. Three-dimensional sketch showing the major environments for glacimarine sedimentation in a cool-temperate climate, based on the present north-northwest Gulf of Alaska (not true orthographic scale). Numbered features are as follows: 1 - - Tidewater front (see Figs. 4 and 7 for details). 2 - - Side drainage with gravel pocket beach and talus fan on fjord floor. 3 - Bergstone mud on glacial fjord floor. 4 - - Sandur plain in marine outwash fjord. 5 - Sediment gravity flow channels in marine outwash mud. 6 - E n t r a n c e sill capped with morainal banks and submarine outwash fan. 7 - Continuation of glacial trough onto the shelf. Relict sediment gravity flows on floor and relict till on walls; presently infilled with marine outwash mud. 8 - - Rocky shore with gravel pocket beaches. 9 - - Tectonically uplifted banks of older rocks exposed by winnowing on the shelf. 1 0 - - Extensive tidal mudflat in estuary. 1 1 - - Recurved and cuspate spits from alongshore transport to north-northwest (top left). Relict moraine at estuary entrance formed at glacier margin during periods of lower sea level. 12 -Marine outwash mud on the continental shelf. 1 3 - - Relict sand deposits. 1 4 - - Thick temperate rain forest of spruce/hemlock with muskeg swamps. 1 5 - - Raised marine terrace. 1 6 - - Eolian dunes fed from spit at wave-dominated delta. 17 -Large slide/slump areas on the continental shelf. 1 8 - - Ice-dammed lake with sublacustrine outwash fan and lake laminites with iceberg-rafted debris. 1 9 - Terminal moraine and gravelly beach. 2 0 - Sandur/delta system with small moraines and gravelly beach. 2 1 - - Modern littoral sand. 2 2 - - Bank on continental shelf of relict glacier-marginal sediment, currently being winnowed. 2 3 - - Tidewater glacier just at sea level with short-headed stream, fan deltas. 2 4 - - Riverdominated estuary with stunted barrier island system offshore. 2 5 - - Sea valley on continental shelf being infilled with marine outwash mud.

these distinctive proglacial processes, many others in both proglacial and more distal paraglacial settings are similar, and for brevity they are often considered together in this discussion.

Fjord systems Marginal deposits The area in contact with the glacier margin at present tidewater fronts includes morainal bank and submarine outwash systems. Morainal bank systems can consist of several

different facies because of the variety of processes that contribute sediment at a grounding line (Fig.4). The term morainal bank is preferred over other terms such as moraine because it differentiates between subaerial and subaqueous deposition and their very different processes of formation (Powell, 1981, 1983b 1984a). Morainal banks form when a terminus is at a stable or quasi-stable site and can be associated intimately with submarine outwash. The length of time a glacier front remains at one location will control the ultimate size of a

:~70 MELTOUT

X

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Fig.4. Processes and lithofacies associations contributing to morainal bank systems at tidewater fronts of temperate glaciers.

morainal bank as lodgement, meltout, dumping, and push and squeeze processes contribute debris from the glacier (Fig.4). The following example serves to illustrate this point using only the meltout criterion, but it provides some concepts as to how debris may be distributed at a tidewater front. At McBride Glacier, Glacier Bay, ocean currents on average flow at about 0.03 m s ~-1 below a 3-10 m water depth and water near the glacier is diluted by direct melting of the glacier face (Cowan and Powell, 1986). At those depths, water temperature annually averages between 2 ° and 4°C. Weeks and Campbell (1973) presented a mathematical expression for glacier ice melting by seawater. Using their equation for McBride Glacier with the following variables: temperature difference between ice and sea = 2°C, water velocities = 0.03 m s--l, and the height of the glacier in contact with water being about 200 m, the rate of ice loss from back-melting by seawater is 8.9 m a-1. Combining t hat melting rate with the volume of basal debris calculated above, 13.4 m 2 a-1 of basal debris can be melted out per

meter length of grounding line. If the frontal position was quasi-stable then t h a t would contribute to building a morainal bank. Measured frontal r e t r e a t rates of McBride Glacier are 250 m a- 1 at present. Therefore, instead of building a morainal bank, the basal meltout debris is spread out as a layer about 0.05 m thick, per meter length of grounding line. Volumes of these deposits will vary among different glaciers and at different times during r e t r e a t as the variables change. Other estimates of sediment accumulation rates near a subglacial stream outlet at McBride grounding line indicate a maximum of 13 m a -1 of sediment accumulates (E.A, Cowan, pers. commun., 1987). Therefore, the relative thickness of debris produced per year at the retreating grounding line is: 0.05 m (0.4%) basal melting, 3.95 m (30.6%) stream bedload and mass-flow deposits, and 9 m (69°//0) suspension settling. Other processes such as glacier push, subglacial lodgement, submarine subglacial stream transport, and others shown in Fig.4, can contribute furt her to morainal bank growth (Powell, 1981, 1983b). In fact, some of the

371 temperate-margin morainal banks of Pleistocene age in Maine can be made up of large proportions of submarine outwash deposits (Smith, 1982; Thompson, 1982), commonly for tens of kilometers of strike along a grounding line. These deposits indicate t hat the margin was sometimes lifted off its base, perhaps by tidal pumping, which allowed subglacial meltwater to flow out to the morainal bank, probably as sheet flow. These examples illustrate t hat morainal bank systems are complex and vary from one site to a n o t h e r depending on glacial regime and dominant processes. Surface grab samples of modern morainal banks include diamicton in pockets between poorly sorted sandy gravel (to cobble size) or gravelly sand (Powell, 1981, 1983b). These observations have been confirmed by bottom video filming but it has also shown the largest particles are boulders. Bank sediments are resedimented by slides and slumps and sediment gravity flows down the fore-slope of the bank (Powell, 1983b; Carlson, this issue). They may also collapse on their up-

glacier side once the support of the glacier is lost during and after r e t r e a t (Powell, 1981, 1983b; Seramur and Powell, 1987). A morainal bank has been observed to form at the face of McBride Glacier between 1978 and 1981. Since then, the tidewater front has retreated a further 1 km (Fig.5); the bank it built did not create stability for the front, probably because the glacier continued to be under-nourished. The front had quasi-stabilized previously at the present entrance to McBride Inlet for about 30 years (1946-1976). On both occasions of quasi-stability, the morainal bank systems have had contributions from major subglacial streams and contain large volumes of submarine outwash. The 1978-1980 morainal bank at McBride Glacier is intertidal, much of its surface being exposed at spring ebb tides. Its total volume approximates 1.2 x 106 m 3 ( a rate of 4 x 105 m 3 a - l ) , and its form is asymmetric (Fig.5). It is highest at its nort h end where boulders are exposed on its surface and to the south the surface progressively changes from sand to

Fig.5. McBride Inlet, Glacier Bay, showing the entrance morainal bank made evident by its last stages of sedimentation, subaerial deltas, and a younger intertidal morainal bank (betweenthe 1978-1980glacier fronts) that is cappedby a submarine outwash lithofacies association. The 1986 coastline and bathymetry have a mean lowest low-water datum.

;~7S mud (Fig.5) and is littered with pebbles, cobbles and boulders. On the down-glacier side the subtidal fore-slope is 15 ~:' ( 6 on the intertidal surface), whereas up-glacier the subtidal leeslope is 9 '~ (4" on the intersurface). Sediment exposed at the top of the morainal bank is submarine outwash with contributions of supraglacial and englacial debris dropped at the tidewater front by meltout and calve-dumping. Additional sediment is contributed today from icebergs th at float over or become stranded on the bank. Imbricate boulders at the north end of the bank are th o u g h t to be the last deposits of a subglacial stream, having been deposited below sea level right at the stream efflux. Fewer boulders were transported downstream on the fore-slopes on the bank (Fig.5). Some are angular and were dropped by calving icebergs from a supraglacial position. The na t ur e of the sediment surface is Very similar to that observed in up to 60 m of water at tidewater fronts from bottom video cameras. To the south, the bank is made mostly of sand and is capped by lenses of mud from recent intertidal deposition and ice-rafted gravel and rubble from grounded or floating icebergs. The mud can be stratified with horizontal sand laminae from tidal fluctuations. Elsewhere the mud is contorted or homogeneous with mixed sand and gravel (a diamicton) due to iceberg turbation. A pit 1.6 m deep, dug in the sandy area, is almost entirely muddy, very coarse sand or granule gravel which is very crudely bedded on a 0.5-- 1 m scale, and where contacts are indistinct and amalgamated. Within these are beds 0.2-0.5 m thick of structureless, moderately sorted, fine to medium sand; structureless muddy sand t ha t can have rare laminae of very fine sand; and thin to medium beds of well-washed pebble gravel with rounded clasts (Fig.6). The muddy sand and gravel, because of poor sorting and crude bedding with amalgamated contacts, are interpreted as sediment dropped from suspension as the turbid freshwater plume lost its competency as it rose buoyantly or continued as an underflow beyond its efflux. The

better sorted sediment was deposited at higher discharges when the stream was in full traction phase at the site and the small lenses of! finer grained sediment are rare remnants of sediment that accumulated at the site probably when the discharge jet was directed in a n o t h e r direction. Submarine discharges from subglacial streams are extremely important for glacimarine sedimentation for two reasons. First, they contribute to grounding-line sedimentation by producing submarine outwash (Fig.7). Secondly, in this temperate setting they contribute by far the greatest proportion of glacial rock flour of which much glacimarine sediment is composed. Active submarine outwash fans are extremely difficult to sample because of their coarse particle size and because of icebergs calving from the front and floating in the sea. It has been possible to sample water salinity, temperature and sediment concentration at an upwelling from a submarine, subglacial discharge, and sediment traps have been deployed to determine sedimentation rates (Cowan and Powell, 1986). Attempts to monitor sea-floor processes have thus far been thwarted by very rapid sediment accumulation rates, sea-floor instability and icebergs destroying sampling operations or equipment. However, recent subtidal deposits have been described above and some remote sensing techniques have allowed information to be gathered. Sediment accumulation rates are so high that it is possible to conduct echosounding surveys in successive years to determine net accumulation by bathymetric changes. From the surveys, submarine outwash was found to accumulate at a rate of 5 m a ' at the efflux of a subglacial discharge from McBride Glacier. Furthermore, these deposits have been recognized in seismic reflection profiles by detailed seismic stratigraphic analyses, at sites where historic tidewater-front positions have been mapped (Seramur and Powell, 1987). They are intimately associated with what are interpreted as morainal banks, and have fan- or wedge-shaped geometries most commonly extending from a bank. They very likely consist

I

37.~

Fig.6. A modern morainal bank exposed at low tide (a) capped by at least 1.6 m of submarine outwash (b) which is mainly muddy very coarse sand or granule gravel with scattered boulders and indistinct, welded contacts. The bank surface shows an up-glacier gravel lag (c) with recently iceberg-rafted angular gravel (d). Other areas of the bank have intertidal muds and iceberg wallows (e).

of coarse-grained sediment based on the conc e n t r a t i o n of h y p e r b o l i c seismic reflectors in proximal areas. F u r t h e r m o r e , particles as coarse as medium sand are trapped in suspended sediment traps: h a v i n g settled from

turbid overflows at an upwelling from a submarine, subglacial discharge; m u c h coarser bedload is dumped at the efflux. Thus far, this section on glacial m a r g i n a l deposits has discussed m a r i n e systems. How-

375

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X GR/~VII"Y" F - I . ~ V S

UNDEHFLUW5

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Fig.7. Processes and lithofacies associations of submarine outwash produced from subglaeial discharges at tidewater fronts of temperate glaciers.

ever, if glacial flux is such that the base of the terminus can be maintained at sea level, transitional systems are produced in contact with the grounded ice cliff. In fjords such as at Reid and Riggs glaciers, Glacier Bay, glacial contact debris can remain relatively unmodified by wave action even in shallow water, and morainal banks can form. Due to relatively quiet water, mud can also accumulate along a margin away from stream activity. Fluvial discharges are most commonly at or near sea level in this shallow water and they form fan deltas that prograde across the glacier front and down-fjord. Once these deltas aggrade, they provide stability for the glacier margin and iceberg calving may cease. If so, melting takes over as the major ablation process and the glacier snout obtains a standard terrestrial hyperbolic shape to its axial profile.

Proglacial and paraglacial deposits Proglacial and paraglacial marine systems occur in, respectively, glacial and marine

outwash fjords. In glacial fjords, the two major direct contributors of glacially derived sediment to the sea are meltwater streams and icebergs (correctly, "growlers" or "bergy bits"). Proglacial sediments proximal to tidewater fronts are commonly interstratified on a bed or lamina scale where sand or silt layers occur between mud layers. In sedimentary rocks, these are termed laminites (cf. Edwards, 1987). Major sources that produce interstratification by episodic depositional or erosional events are: (1) stream discharges, (2) tidal currents, (3) wind-generated waves and currents, and (4) subaqueous slumps and sediment gravity flows. Variation in stream discharges commonly produce cyclopsams (sand-mud couplets) and cyclopels (silt-mud couplets) (Fig.6). These sediments have been sampled both as modern facies on the sea floor and in suspended sediment traps at different depths through the water column (Cowan and Powell, 1986; Cowen et al., 1987). They originate by settling from turbid overflow plumes, at a maximum re-

37(;

corded rate of 15.4 cm dry sediment in 19 h (Cowan et al., 1987), all of which was of sand size (cyclopsams). During that period, marine interflows also occurred because of the extremely high suspended sediment concentrations which were undoubtedly caused by a rainstorm event. These sediments can be preserved in the rock record, because they were first described and defined in long cores (Mackiewicz et al., 1984) and are recognizable in ancient successions (e.g., Mustard and Donaldson, 1987; Powell, 1987). The distribution of cyclopsams and cyclopels and the number of laminae produced per day are controlled by variations in meltwater stream discharge, tides and wind shear. Each of these controls is independent in timing and magnitude and therefore, it is extremely difficult to predict the number of laminae but many can be produced in 24 h (Cowan et al., 1988). Cyctopsams generally occur within 1 km of a grounding line and are intimately associated with coarser submarine outwash. Cyclopels may be deposited proximally either during low discharges or along the tidewater front away from the submarine stream efflux. They are more common between one and several kilometers from a tidewater front because they are less diluted by other deposits (Mackiewicz et al., 1984). With greater distance from the submarine efflux, they are lost because all individual silt particles have settled and sedimentation rates may have decreased sufficiently for bioturbation to become significant. Slumps and subaqueous sediment gravity flows may be generated by (1) biogenic gas, (2) glacier push from summer-winter glacier-front fluctuations and from "jerky motion", (3) earthquakes, (4) high sedimentation rates creating slope over-steepening and entrapment of porewater, or redistribution from the mouths of subglacial streams, (5) wave-induced liquefaction, (6) tidal forcing causing liquefaction, and (7) iceberg calving. The first three processes occur much less frequently and less regularly than the last four, which, given fluctuations in sedimentation rates, could result in variations in the degree of stratification

development. Although the nature of' sediment gravity flow deposits cannot be used to identify the cause of resedimentation, iceberg calving and slope over-steepening are probably the most common causes at temperate tidewater fronts. A mathematical model has been formulated which predicts that an iceberg calved above water level from the face of a tidewater front can either impact the floor or come sufficiently close, so that it or its preceding pressure wave could redistribute bottom sediment by sediment gravity flows (Powell, 1985). Iceberg calving from tidewater fronts is implied from the modelling to be a very important process for generating ice-proximal sand and/or silt laminae and/or beds in bergstone muds or bergstone diamictons. Beyond the interstratified sediments, and also interbedded with them, are homogeneous muds that are marine outwash deposits of glacial rock flour. In glacial fjords these muds commonly contain iceberg-rafted debris (IBRD) and are termed pebbly mud (or lonestone mud; non-genetic facies terms) or dropstone mud or bergstone mud when genesis is known and emphasized (Powell, 1984a). They can grade in to and out of diamicton (> 10% pebble sizes; Powell, 1984a) depending on relative rates of iceberg rafting and marine outwash production. Ice turbates (Reimnitz and Barnes, 1974) are rare in the fjords because once icebergs are floating after the calving event, they are too small to touch bottom unless they are in very shallow water. Muds can be interbedded with interstratified facies if paths of turbid overflow plumes vary and are probably commonly formed over a proglacial basin during winter when iceberg calving still occurs but when stream discharges decrease drastically (although they do not cease entirely) and cyclopels and cyclopsams do not form (Cowan and Powell, 1986). Accumulation rates of bergstone mud appear to decrease exponentially from a tidewater front with the proportion of IBRD depending greatly on the paths of the icebergs. Depending on oceanic circulation, strengths of meltwater

377 discharge and katabatiac winds, sea ice, and sea-floor topography, icebergs may either stay close to a tidewater front or move away at different rates. The mixing rates of mud and IBRD determines the end product. For example, diamicton may form proximally to a glacier front if all bergs remain close to it, as can occur. However, if icebergs move out away from the face, diamictons may also form distally because mud accumulation rates are lower. The calving speed of McBride Glacier can be estimated by simple mass balance: A S = S , - So - S ~

where these are speeds related to frontal change ( A S = - 2 5 0 m a - l , an observation of retreat), glacier flow (S~=300 m a -1, determined using aerial photographs for surface flow, of which about 95% is probably basal flow (cf. Kamb et al., 1985)), glacier melting (Sin = 10 m a-1, an average from calculations determined above), and calving speed (So). The calving speed is 540 m a-1 and given total basal debris thickness of 1.5 m and a basal glacier width of 500 m, 4x105 m a a -1 of sediment is ice rafted into McBride Inlet. Because of a narrow entrance, most icebergs are trapped in the inlet and lose their debris locally; consequently, comparative sediment accumulation rates can be made. Average sediment accumulation rates of mud over the entire inlet (area about 7 x 105 m 2) is about 2 m a - 1 (E.A. Cowan, pers. commun., 1987) giving a total volume of mud of about 1.4 x 106 m 3 a - 1. If average accumulation rates are assumed to be evenly spread throughout the inlet, then about 22% of debris is iceberg rafted and 78% is from meltwater; a proportion that is approximated in cores of bottom sediment, in spite of a wide variation in the natural system. Marine outwash fjords contain little or no IBRD within ubiquitous marine outwash mud (Powell, 1981, 1983b, 1984b). In addition to an absence of IBRD, the mud is characterized by black horizons as in Queen Inlet (Hoskin and Burrell, 1972) or yellow-gray layers (Powell,

1983b). Black layers show no enhancement of organic carbon (Hoskin and Burrell, 1972; Loder and Hood, 1972) and although their origin is vague, they are thought to be annual winter layers (Hoskin and Burrell, 1972; Hoskin et al., 1976). Yellow-gray layers are considered a result of variation of redox conditions in the sediment column (cf. Carsola, 1954; Turner, 1971). Redox variations may be due to changes in sedimentation rates or be caused by mud having different sources: blue-gray mud being a direct glacial input, yellow-gray mud having been reworked from an oxidizing intertidal zone (Powell, 1983b). Heads of marine outwash fjords are dominated by sandur plains and deltas, or tidal flats. Seismic reflection profiles show channels on the floor of Queen Inlet, Glacier Bay, originating at the fjord-head delta. Seismic reflection profiles show the channels have depositional-erosional forms exhibiting levees and terraces (Powell, 1983b; Carlson, this issue). Four branching channels cut the floor; the main channel is 10.5 km long and up to 259 m wide, and has an eroded depth of 32 m and an average slope downfjord of 1.3 ° (Carlson et al., in prep.). The channels do not necessarily follow buried topography and were relatively recently eroded into fjord-bottom sediment. Channel-floor sediments are much more sandy than on the rest of the fjord floor (Hoskin and Burrell, 1972). These sediments are rhythmically layered sand and mud (Carlson et al., in prep.) that is thought to originate from sediment gravity flows from the delta front (Hoskin and Burrell, 1972; cf. Prior et al., 1986). Another marine outwash fjord that has been well studied is Cook Inlet, whose features have been summarized by Rappeport (1980). He defines six depositional zones: (1) high-energy shoreface, (2) trough-edge platform, (3) trough slope, (4) trough floor, (5) trough-mouth plateau, and (6) seaward progradational ramp. Of relevance here is the trough-edge platform which extends 10-15 km offshore to about 60 m deep, and has gravel, pebbly sands, patchy

37S clean sand, shell hash, and some silt and clay. The coarse-grained sediment is due to tidal currents that may flow at 30 60 cm s The trough slope is mantled by coarse gravel and cobble (glacial lags) with sporadic thin sand ribbons and patches created by tidal currents flowing up to 40 cm s 1. Sand waves higher than 7 m dominate the trough floor where current speeds reach 20 cm s 1 The sand waves are often covered with ripples, and they migrate over a highly bioturbated sandy surface that has abundant sand dollars, sea pens, and concentrations of fine-grained organic materials. Organic carbon concentrations are low in glacial ice and meltwater streams (Loder and Hood, 1972). They are highest in fjord waters during summer, just beneath the surface mixing layer, but most terrestrial organic matter is carried out of fjords in suspension as reflected in bottom sediments which have a very low organic carbon content. In shallow waters and offshore banks, macrofaunal communities are quite abundant and include barnacles, echinoids, bryozoans, pelecypods, gastropods, sponges, algae and biogenic debris (e.g., Hoskin and Nelson, 1969, 1971). Temperate-water corals also grow in the region (Sharma, 1979). The other major biogenic component is phytoplankton; two prolific seasonal blooms of diatoms occur each year, which is reflected in the silica content of bottom sediments (Sharma, 1979). Numbers of diatom tests in bottom sediments are most commonly low because of the higher rate of production of terrigenous components.

Continental shelf systems Marginal and proglacial deposits As recently as 70-80 years ago, some areas of the continental shelf of the Gulf of Alaska were proglacial because large numbers of dirty icebergs entered the gulf (Molnia and Carlson, 1980). The only current example of proglacial systems on the continental shelf is the La Perouse Glacier whose terminus just ends in

the sea as a cliff. At the grounding line, basal debris is melted out and supraglacial debris is dumped during calving. These sediments are quickly reworked by wave action along the storm-dominated Gulf of Alaska shelf, leaving a gravel foreshore with typical beach-face deposits that contain some extremely large boulders whose presence are good indicators of glacial transport. Subglacial and marginal streams are shortheaded when they discharge into the sea. Those that flow directly from the glacier into the sea contribute to the gravelly and sandy beach deposits. Marginal deltas are wave dominated, their sediment being quickly sorted and, in part, reworked laterally into beaches. Mud, originating as glacial rock flour, is transported offshore mainly in turbid surface plumes. Sedimentary processes and facies under other proglacial conditions have to be inferred by using: (1) the excellent 3000 m-thick glacimarine succession of the Yakataga Formation that extends in age from the Early Pleistocene to about 6.3 Ma (Armentrout, 1983), and (2) a few relict facies on the continental shelf combined with predictions from information known about modern fjords. The Yakataga Formation is made up of interstratified matrix-supported diamictites, conglomerates, sandstones, mudstones, and laminated siltstones (Plafker and Addicott, 1976; Armentrout, 1983; Eyles, 1987, 1988). They are interpreted as upper bathyal to shelfedge deposits (Lagoe, 1983; Eyles, 1988) and, within intraformational megachannels, as fjord- (Armentrout, 1983) and shelf/slopetrough (Eyles, 1987) deposits. Water masses at the time were cold-subarctic to cool-temperate (Lagoe, 1983). This is an excellent record for providing glacial geological information about the south-southeast Alaskan region; however, there is no land record of the Upper Pleistocene and Holocene in this succession. The Upper Pleistocene continental shelf record has to be obtained from marine geological studies, for which comprehensive data are currently lacking.

379 While on the continental shelf, Quaternary glaciers are known to have occupied submarine valleys which have diamictons on their walls. The diamictons and other features have led Carlson et al. (1982) to suggest that these valleys are structurally controlled, but glacially eroded. During Wisconsinan time, the Alaskan Cordilleran Ice Sheet was relatively small (Hamilton and Thorson, 1983) so even at glacial maximum the Gulf of Alaska shelf may not have been as deep as present-day Antarctica. Consequently, much of the shelf may have stood above sea level and some glaciers may not have reached the sea (Goldthwait, 1986; Mann, 1986b). Those glaciers that did reach the sea probably ended as tidewater cliffs even during past glacial maxima. Ice shelves must have been absent because the 0 ° isotherm was above sea level in the gulf (CLIMAP, 1976) and refugia were present along the coast (e.g., Mann, 1986b). Only very short and narrow floating margins may have rarely formed from ice streams where ice flux was much greater than calving speed. The major environmental difference beyond tidewater fronts in fjords and on a shelf is wave climate. Present storm waves reach 40 m high, to be an easy influence on most shelf sediment (Molnia, 1987). Large storm waves should produce sedimentary structures such as hummocky cross stratification and swaley cross stratification when there is a better sand supply than today. Such features have been found in intertidal zones under modern conditions (Bartsch-Winkler and Schmoll, 1984). Storm waves can redistribute sediment in geostrophic flows or turbidity currents (Nelson, 1982; Walker, 1985), and any bank areas can be effectively winnowed. Relict and palimpsest facies on the outer shelf and on platform and bank areas (Carlson et al., 1977), such as Tarr Bank and Fairweather Ground, were probably glacial-contact facies which have been winnowed by waves and bottom currents (Carlson, this issue). The only confirmed morainal banks exposed on the sea floor at present are those at the mouths of fjords composed of sandy or gravelly clayey silt.

Ridges offshore from Icy B a y - Y a k u t a t may be morainal banks but are probably terrestrial moraines deposited during retreat of the last glaciers at times of lower sea level (Molnia, 1986). The best preservation of all primary glacimarine facies, especially those defined by more delicate structures, will occur below storm wave base. BiOlogical activity will then be an important consideration, as indeed it would be over the rest of the shelf too. Occurrence and number of sediment gravity flow deposits depends on local relief and sedimentation rates on the shelf and proximity to the grounding line.

Distal paraglacial deposits The present continental shelf and slope of the Gulf of Alaska is mostly a paraglacial system receiving its sediment from fluvial discharges on land. Most available information is from the Y a k a t a g a - Y a k u t a t area (Wright and Sharma, 1969; Carlson et al., 1977; Molnia and Carlson, 1978; 1980; Armentrout, 1980). In general, Holocene shelf sediment fines offshore from littoral sand (to a depth of 30 m, 7 km offshore), to silty sand (to a depth of 60 m, 20 km offshore), to clayey silts (at least to a depth of 203 m, 75 km offshore) (Armentrout, 1980). The clayey silt is accumulating at average rates of 15 mm a-1 (Carlson et al., 1977) and the total volume of Holocene sediment is estimated at 3 x 103 km 3, which gives an average thickness, if spread over the entire shelf, of 55 m (Molnia et al., 1978). The Alaskan Coastal Current redistributes sediment and most Holocene sediment is assumed to have been contributed by existing drainage (Molnia et al., 1978) with little by-passing the shelf. High areas on the shelf apart from the relict glacial contact systems are shown in seismic reflection profiles to be exposed folded Tertiary bedrock (Molnia and Carlson, 1978). These are old tectonostratigraphic terranes and uplifted subduction complex rocks that are also kept exposed by wave winnowing. Depressions on the shelf, such as the submarine valleys, are partially filled with Holocene clayey silt,

380

similar to that on the rest of the shelf, indicating they trap suspended sediment. Slumping from their heads redistributes sediment (Carlson et al., 1982), but only locally such as west of Kayak Island, is near-bottom sediment known to move beyond the outer shelf. Slides and slumps are common features of sediments on the north-northeast shelf and fjords of the Gulf of Alaska (e.g., Carlson and Molnia, 1978; Carlson, 1978, Schwab and Lee, 1983). Some of the features are quite large (areas up to 1080 km 2) and they often have scarps at their head, thus appearing to have propagated by headward growth. The marine sediments are underconsolidated (Hampton et al., 1987), and when shaken by large earthquakes that occur in the area, they are thought to move en masse (Carlson, 1978). Movement also may be initiated often to a greater degree by intense storm activity (Schwab and Lee, 1983). Biofacies are important aspects of the glacimarine environment to include in sedimentary models but details are beyond the scope of this discussion. Distribution of modern plankton in the Gulf of Alaska region as well as the productivity and community compositions have been recently summarized by Sambrotto and Lorenzen (1987) and Cooney (1987). Benthic foraminiferal and mollusk distribution on the south-southeast Alaska continental shelf have been recently summarized and related to types of substrate and water depth (Echols and Armentrout, 1980; Hickman and Nesbitt, 1980; Quinterno et al., 1980). Very little surface sediment sampling of the continental slope has been carried out. Parts of the upper continental slope are covered by the same clayey silt present on the shelf (Carlson et al., 1977; Molnia and Carlson, 1978; Armentrout, 1980). However, in other areas where bottom currents are thought to be strong enough to prevent sediment deposition today, and Quaternary greenish gray pebbly mud is present from glacial advance over the shelf (Carlson et al., 1977; Molnia and Carlson, 1978). The Alaska Current impinges at the

shelLslope break but very little is known about its effect on bottom lithofacies.

Fluvial-deltaic systems Proglacial fan deltas have plains consisting of coarse-grained sediment because of the proximity of the glacial source. Along the Gulf of Alaska coast, sandur deltas are produced by the Bering, Grand Plateau and Finger Glacier and most of Malaspina Glacier. Of these outwash plains that end as deltas, the fullest description is of the Yana Stream that has built a sandur from Malaspina Glacier (Gustavson, 1974; Boothroyd and Ashley, 1975). Facies of the plains are gravelly longitudinal bars with minor large-scale cross-bedded sands on the upper section, flat-bedded and largescale cross-bedded sands and sandy gravels in the middle, and large-scale planar and trough cross-bedded and ripple-drift cross-laminated sands on the lower section at the delta. The coastline areas are mainly wave dominated where mouth bars and spits extend several hundreds of meters northwest-west due to the prevailing wave climate and alongshelf currents. Inactive areas of the delta at the shoreline can be reworked into beach ridges with lagoons behind. Prodelta sediments are sandy to depths of about 50 m and up to 15 km offshore; clayey silt then extends out to the upper continental slope (Armentrout, 1980). The offshore areas here, as well as offshore from Bering Glacier are sites of major sediment slides and slumps (Carlson and Molnia, 1978; Carlson, 1978; Schwab and Lee, 1983). The major paraglacial delta on the gulf coast, fed mainly from runoff, is a strongly river-influenced delta building in a small estuary at the Copper River. The sediment is well-sorted medium to fine sand. Directly offshore is a stunted mesotidal barrier-island system and most sand is transported northwest as alongshore drift, although there is an offshore component carrying rifler sediment (Sharma, 1979). The Copper River discharge is 5% of the regional total (Royer, 1982) and is the major sediment contributor for the

381 northern gulf, and discharge peaks reach 7×103 m 3 s -1 during summer. Suspended sediment concentrations are up to 1700 mg l- 1, with a total sediment production of about 1.07 × 1011 kg a -1, (Reimnitz, 1966), of which at least 15% is bedload. Molnia et al. (1978) estimated that the Copper and Malaspina drainage basins combined produced 0.15 km 3 a-1 of glacial rock flour to the gulf, which is about 30-50% of the total input. The Copper River system and its associated barrier island have been described sedimentologically by Galloway (1976). The subaerial delta plain facies are marsh and swamp organic muds, and braided to estuarine distributary channel fills. The tidal lagoon is composed of tidal sand and mudflat sequences interlaced with a complex of tidal channel fills. The shoreface consists of marginal island, breaker bar, and middle shoreface sands, lower shoreface sand and mud, and prodelta/shelf mud. Deltas in fjords can be fan deltas when fed by large fluvial discharges, but in contrast to those on the open shelf coastline they can also be tide dominated. At the former, delta plains are dominated by braided streams, with gravel bars and sandy sheet-flood plains. Macrotidal inlets produce proglacial, stream-dominated deltas where delta plains have extensive tidal flats cut by coarse braided channels of the meltwater stream. Some unique processes occur within these proglacial systems under strong tidal influence. Facies of fan deltas are dominated by meltwater streams whose hydrographs frequently show diurnal temperature and melting fluctuations on which rainfall events and tidal prism changes are superimposed (N.D. Smith, pers. commun., 1987). These deltas can form very rapidly and prograde and aggrade in front of the tidewater cliff just as a glacier retreats from the sea. The deltas develop from submarine outwash fans when a glacier front becomes quasi-stable. A delta in a fjord system has been observed to form during retreat of Riggs Glacier, Glacier Bay (Powell and Cowan, 1986). In 1979, the glacier terminated in about 55 m of water (Powell, 1983b); by 1981, a delta

with a fully intertidal plain had prograded 200 m from the glacier. By 1985, the delta plain had aggraded above tidal range and prograded 100 m farther into what had been 32 m of water in 1981 (Powell and Cowan, 1986); a total of about 106 m 3 of sediment in 4 years or 0.25 × 106 m 3 a - 1 had accumulated. In the early stages of delta formation when the system is transforming from a submarine outwash fan into a normal delta, the plain is fully intertidal and stream channels change position with each tidal cycle. During flood tide, the stream is buoyantly lifted off the delta plain; then with ebb tide, the stream is progressively lowered back to the delta plain and can erode completely new channels each time. Consequently, very little sediment is sorted on the delta plain and aggradation occurs only by changing grade due to progradation. Resedimentation events from the delta plain are, therefore, very common and prodeltaic sediment is stratified sand and mud. A couplet can be generated over a tidal cycle (Powell, 1981, 1983b). As the delta plain aggrades, channels become quasi-stable and sedimentation is influenced by sediment discharge from subglacial streams as well as by tides. Pulses of high sediment concentration are generated during peak discharges, but major events for prodeltaic deposition occur during tidal draw-downs over monthly spring tide periods where low tide level is at or near the break between delta plain and delta front. Then, large quantities of sediment (up to 44 g 1-1 in suspension; Woodland et al., 1987) temporarily stored on the delta plain are transported down the front into prodelta areas (Smith et al., 1986). Resedimentation by mass flows are also common from the delta front (Phillips et al., 1987). Mass movements result in delta front and prodelta facies that are stratified, where the bi-monthly sandy mass-flow deposits produced by tidal draw-down and are probably thicker than resedimented units produced by failures from oversteepening on the delta front. Paraglacial deltas fed mainly by runoff have not been described from fjords in the region,

but have been studied recently at other locations (papers in Syvitski and Skei, 1983). Major differences between proglacial and the more distal paraglacial types are t ha t the latter have less "flashy" discharges because of vegetated runoff basins, and they have a much greater organic component (both floral and faunal). The delta plains are mainly braided stream systems with channels ending at ephemeral distributary mouth bars (cf. Syvitski and Farrow, 1983; Kostaschuk, 1985). Tidal flat sedimentation is common in intertidal areas between distributary channels. Delta front and prodeltaic areas are sites of common redepositional processes: slides and slumps can occur off the delta front (cf. Prior et al., 1981, 1982) and channels incised in delta front and prodelta areas sometimes extend many kilometers down-fjord (cf. Prior et al., 1986).

Beach systems Beaches in fjords are the only ones in the region th at have a glacial signature except for the glacial contact beach at La Perouse Glacier. Small bergy bits can be stranded on fjord beaches after floating to shore during flood tide. These icebergs may remain stranded until they melt completely, especially if they are stranded during spring tides. Piles of icerafted debris are added to intertidal environments by this process. DistinctLve debris includes gravel-size clasts, till pellets (Ovenshine, 1970) and mud clasts. The latter are

derived from englacial and basal positions in the glacier (Heiny and Powell, 1982), and, like till pellets, are frozen but they lack coarser particles. Gravel-size IBRD in intertidal zones also provide sites for encrusting organisms. "Iceberg rosettes" are formed by stranded icebergs in intertidal zones on sandy shores (Fig.8). An iceberg melts totally in situ and shore sand adheres to its surface by meniscal water tension. With progressive melting, successive ridges several millimeters high and several grains thick form in concentric patterns as the ice dimensions decrease (cf. Rubin and H u n t e r (1984) for similar structures formed by hailstones). They may be draped with mud in some intertidal zones. Larger icebergs are rocked by waves and tidal currents to form iceberg wallows (cf. Dionne, 1976; Barnes et al., 1982; Reimnitz and Kempema, 1982). If icebergs do not ground completely but move in shallow water with tidal currents or by wind, then their keels can produce shallow scour marks. These can take various forms, but grooves (smooth and chevron), chatter, bounce, and roll marks are common (Fig.9), and may be preserved as bedding-plane structures. Along the gulf coast, beach systems vary in texture as a result of different sediment sources from local fluvial or alongshore transport, and wave climate. Gravel beaches may occur as pocket beaches, concent rat ed by local erosion of bedrock or moraines, or in areas downdrift from actively eroding glacier mar-

Fig.8. Iceberg rosettes in the process of forming (a) and the final sedimentary structure (b).

383

Fig.9. Iceberg scour marks on beach faces: Smooth groove (a), and chevron groove and c h a t t e r mark (b).

3~4

gins (cf. Molnia and Wheeler, 1978; Hayes, 1986). Sand and mixed sand and gravel beaches occur downdrift of outwash streams. Spits are commonly constructed across by mouths (recurved), at narrow restricted passage ways (cuspate), or at river mouths (wave-dominated delta forms), all due to alongshore drift of sediment both on the gulf coast and inland waterways. Back-beach lagoons are common features at bayhead beaches, bay mouths, and river mouths. They are commonly highly organic, and are often inundated by washover fans from frequent storm action. On drier back-beach areas, but also within storm-surge range, vegetated meadows and/or forests are common. Local eolian dune systems occur where sediment is available from spit, beach, and washover deposits. Extensive salt marshes can be present locally in estuaries and fjords. All of these features of beaches have been described in a geomorphic sense (Molnia and Wheeler, 1978; Hayes, 1986) but virtually no detailed sedimentological work especially in regard to facies models has been conducted. At present, we assume that these systems being constructed from glacigenic sediment have features similar to those described in the literature from locations beyond glacial influence.

Tidal fiat systems Tidal flats in fjords receive a glacial signature from icebergs in ways similar to those described for beaches. The majority of sediment is fine sand and/or mud that is flaser to lenticular bedded. They lack major biological communities; however, infauna, and floral and faunal encrustions are present. These systems have been briefly described (Powell, 1981, 1983b) but not studied in detail for lithofacies analysis. Along the gulf coast, most tidal flats are distal paraglacial systems. The tidal flat in Knik Arm of upper Cook Inlet is the best described example (Bartsch-Winkler and Ovenshine, 1984; Bartsch-Winkler and Schmoll,

1984). There, a large tidal bore travels 4 m s over medium to fine sandy tidal flats which noticeably lack gravel. The flats have few major tidal channels, but tidal bars are major features. Important sedimentary structures are straight-crested to slightly sinuous ripples, lunate and linguoid ripples and ladder ripples of wind-wave origin. The flats have four main sedimentological zones: high flats of parallel-laminated silt, subordinate cross lamination, plant debris and bubble-cavity sand; intermediate flats with cross-laminated sand and fewer parallel-laminated and bubble-cavity sands; a transition zone of parallel-laminated sand with subordinate herringbone, climbing, and non-climbing cross lamination; and lower flats with parallellaminated sand. These produce an upwardfining sequence analogous to point-bar successions. The deposits are commonly contorted, but low infaunal biological activity means they are not significantly destroyed by bioturbation. Ice floes, which form from December to April, have a significant effect on sediment dispersal because surface sediment freezes to them as they are deposited with each ebb tide. Most of the knowledge of intertidal and subtidal biological communities of the southsoutheast Gulf of Alaska region is from the shelf coastline and less is known about fjord communities. The gulf coast communities have been summarized recently by O'Clair and Zimmerman (1987). Conclusion First-order controls on glacimarine sedimentation in the south-southeast Gulf of Alaska region are: (1) tectonism, which creates accumulation areas for glaciers and provides sources for glacial debris, and (2) climate, which produces high rates of snow accumulation and glacier ablation, and high rates of stream and sediment discharge. Of the second-order controls, glaciers themselves and marine-fluvial interaction are the most important. These valley glaciers end as

385

tidewater fronts that have fluctuated asynchronously among themselves and independent of climate after its initial forcing. Glacial fluctuations are important to glacimarine sediments in two ways. First, depocenters of the highest rates of sediment production migrate with the glacial fronts, producing very timediscordant packages. Secondly, different lithofacies associations are produced under different rates of frontal movement. The cool-temperate to sub-polar climate enhances production of glacial debris. Rates of erosion are high (perhaps about 1.2 cm a-l), glaciers have high debris loads (1.5 m of pure basal debris), glacier flow speeds are high (some over 3000 m a-1), and sediment accumulation rates are high (over 13 m a -1 at a tidewater front). Most glacial debris is ultimately transported to the sea by streams. Fluvial discharges and sediment loads are high both in the gulf and in fjords. Discharges propagate the Alaska Coastal Current which transports most sediment to the present continental shelf. Submarine outwash at tidewater fronts in fjords produce distinctive cyclopsam and cyclopel lithofacies. If a front is quasi-stable, these and other glacial sediments build morainal banks, but if the front moves, then this sediment is spread over the sea floor. Further from a tidewater front, bergstone mud and bergstone diamicton accumulate; in one present-day example 22% of these proglacial sediments is estimated to be from iceberg rafting and 78~/o from plumes of marine outwash. Other depositional systems of the region are fan deltas (that can grow at about 106 m 3 a - 1), wave-dominated beaches along the gulf coast, and tidal flats in bays and fjords. The Gulf of Alaska continental shelf and fjords are dominated by siliciclastic depositional systems with biogenic carbonate material on banks. As an example of a cool-temperate glacimarine basin, it differs markedly from a fully polar setting. The modern Antarctic shelf is characterized by biogenic siliceous sediment and a relative paucity of modern siliciclastic debris.

Although significant research has been conducted along the Gulf of Alaska margin, much more information is required before a comprehensive understanding of the cool-temperate glacimarine environment as a whole is obtained. Of importance are time-series surveys of processes and sediment transport paths and more comprehensive data to establish sediment budgets. Biofacies analysis is badly needed as base-line data for paleoecological models. Compared with many other glaciated margins, the history of glaciation on the southsoutheast coast of Alaska is extremely poorly known. Far more work is required on the shelf and in bays using seismic surveys and long cores to determine glacial fluctuations since Late Pleistocene. An extensive record of earlier glaciations is preserved in the Yakataga Formation.

Acknowledgements The authors are grateful for very helpful comments made on earlier drafts of the manuscript by B.D. Bornhold, P.R. Carlson, C.H. Eyles and J.P.M. Syvitski. A combination of previous work under various grants has been summarized but most recent support is from NSF grants DPP-8420749 and DPP-8619153 to R.D.P.R.D.P. wishes to thank the National Park Service and Exploration Holidays for continued logistical support of field work.

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