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Quaternary Science Reviews 31 (2012) 67e81
Contents lists available at SciVerse ScienceDirect
Quaternary Science Reviews
journal homepage: www.elsevier.com/locate/quascirev
The sequence and timing of large late Pleistocene floods from glacial Lake
Missoula
Michelle A. Hanson a, *, Olav B. Lian b, John J. Clague a
a
b
Department of Earth Sciences, Simon Fraser University, 8888 University Drive, Burnaby, British Columbia, Canada V5A 1S6
Department of Geography, University of the Fraser Valley, 33844 King Road, Abbotsford, British Columbia, Canada V2S 7M8
a r t i c l e i n f o
a b s t r a c t
Article history:
Received 24 March 2011
Received in revised form
28 October 2011
Accepted 10 November 2011
Available online 30 November 2011
Glacial Lake Missoula formed when the Purcell Trench lobe of the Cordilleran ice sheet dammed Clark
Fork River in Montana during the Fraser Glaciation (marine oxygen isotope stage 2). Over a period of
several thousand years, the lake repeatedly filled and drained through its ice dam, and floodwaters
coursed across the landscape in eastern Washington. In this paper, we describe the stratigraphy and
sedimentology of a significant new section of fine-grained glacial Lake Missoula sediment and compare
this section to a similar, previously described sequence of sediments at Ninemile Creek, 26 km to the
northwest. The new exposure, which we informally term the rail line section, is located near Missoula,
Montana, and exposes 29 units, each of which consists of many silt and clay couplets that we interpret to
be varves. The deposits are similar to other fine-grained sediments attributed to glacial Lake Missoula.
Similar varved sediments overlie gravelly flood deposits elsewhere in the glacial Lake Missoula basin.
Each of the 29 units represents a period when the lake was deepening, and all units show evidence for
substantial draining of glacial Lake Missoula that repeatedly exposed the lake floor. The evidence
includes erosion and deformation of glaciolacustrine sediment that we interpret happened during
draining of the lake, desiccation cracks that formed during exposure of the lake bottom, and fluvial sand
deposited as the lake began to refill.
The floods date to between approximately 21.4 and 13.4 cal ka ago based on regional chronological
data. The total number of varves at the rail line and Ninemile sites are, respectively, 732 and 583.
Depending on lake refilling times, each exposure probably records 1350e1500 years of time. We present
three new optical ages from the rail line and Ninemile sites that further limit the age of the floods. These
ages, in calendar years, are 15.1 0.6 ka at the base of the Ninemile exposure, and 14.8 0.7 and
12.6 0.6 ka midway through the rail line exposure. The sediment at the two sections was deposited
during later stages of glacial Lake Missoula, after the largest outburst events.
Ó 2011 Elsevier Ltd. All rights reserved.
Keywords:
Glacial Lake Missoula
Outburst floods
Stratigraphy
Sedimentology
Varves
Optical dating
1. Introduction
The Purcell Trench lobe of the Cordilleran ice sheet filled the
basin now occupied by Lake Pend Oreille, Idaho, during the Fraser
Glaciation (late Wisconsinan; marine oxygen isotope stage 2). The
ice lobe blocked Clark Fork River and impounded glacial Lake
Missoula in the intermontane basins of western Montana (Fig. 1;
Pardee, 1910, 1942). At its maximum, glacial Lake Missoula had
a volume of 2200e2600 km3 and a depth at the ice dam of
610e640 m (Pardee, 1942; Bretz, 1969; Clarke et al., 1984; Craig,
1987; Smith, 2006). When water in glacial Lake Missoula reached
* Corresponding author. Present address: Saskatchewan Geological Survey, Saskatchewan Ministry of Energy and Resources, 200-2101 Scarth Street, Regina,
Saskatchewan, Canada S4P 2H9. Tel.: þ1 306 798 4211; fax: þ1 306 787 1284.
E-mail addresses: mhanson@sfu.ca, Michelle.Hanson@gov.sk.ca (M.A. Hanson),
olav.lian@ufv.ca (O.B. Lian), jclague@sfu.ca (J.J. Clague).
0277-3791/$ e see front matter Ó 2011 Elsevier Ltd. All rights reserved.
doi:10.1016/j.quascirev.2011.11.009
a critical depth, the east end of the ice dam became buoyant,
tunnels opened beneath the dam, and the lake drained catastrophically (Waitt, 1985). The maximum discharge of the largest
flood was at least 17 million m3 s1 (Baker, 1973; Clarke et al., 1984;
O’Connor and Baker, 1992). Floodwaters flowed into glacial Lake
Columbia to the west, depositing a thick layer of gravel, sand, and
silt on top of laminated silty glaciolacustrine sediments (Fig. 1;
Atwater, 1984, 1986, 1987). Floodwaters also flowed across the
Columbia Plateau, eroding Miocene basalt and Pleistocene loess
(Fig. 1; Bretz, 1923, 1925, 1928). The floodwaters formed giant flood
bars and current ripples, deposited fine-grained sediments in some
Columbia River tributaries in southern Washington, and recharged
Columbia River Basalt aquifers with isotopically depleted water
(Bretz, 1925, 1928, 1929; Bretz et al., 1956; Baker, 1973; Waitt, 1980,
1985; Brown et al., 2010). They also backed up southward into
Willamette Valley of western Oregon and deposited slackwater
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M.A. Hanson et al. / Quaternary Science Reviews 31 (2012) 67e81
Fig. 1. Study area showing locations mentioned in the text. The figure shows the maximum extent of the Cordilleran ice sheet as well as the maximum extent of glacial Lake
Columbia. LCFR: Lower Clark Fork River; MV: Mission Valley; NC: Ninemile Creek; RL: Rail line. Darker gray areas were swept by floodwaters from glacial Lake Missoula.
(Figure adapted from Waitt, 1985; Atwater, 1986; Levish, 1997).
sediments there (Allison, 1978; Waitt, 1980, 1985; O’Connor et al.,
2001). Within a few weeks, the floodwaters had reached the
mouth of Columbia River, over 700 km from their source, where
plumes of freshwater spread out into the northeast Pacific Ocean
(Baker, 1973; Craig, 1987; O’Connor and Baker, 1992; Lopes and Mix,
2009; Denlinger and O’Connell, 2010). Turbidity currents triggered
by the floods travelled an additional 1100 km southward along the
floor of the eastern North Pacific Ocean (Zuffa et al., 2000), while
a plume of finer sediment travelled northward along the coast
(Gombiner et al., 2010). Glacial Lake Missoula may have drained
and refilled approximately 90 times during the Fraser Glaciation
(Waitt, 1985; Atwater, 1986).
Most of the evidence for drainings of glacial Lake Missoula has
come from study of flood deposits in Idaho, Washington, and Oregon
(Bretz, 1925, 1928; Bretz et al.,1956; Baker, 1973; Allison, 1978; Waitt,
1980, 1985; Atwater, 1984, 1987; O’Connor et al., 2001). In contrast,
relatively little research has been done in the glacial Lake Missoula
basin itself. Pardee (1910, 1942) described the extent of the lake and
landforms that led him to conclude that very large and rapid currents
were involved in the draining of the lake. Chambers (1971, 1984)
studied a section of cyclic, rhythmically bedded glaciolacustrine
sand, silt, and clay units exposed in a highway cut near Ninemile
Creek, Montana (Fig. 1). Most of the units1 at the Ninemile exposure
comprise a thick layer of very fine sand and silt that grades upward
into rhythmically laminated silt and clay beds that Chambers (1971,
1984) interpreted to be varves. The silteclay couplets thin upward
within units that are separated by unconformable contacts. Thin
beds of pebble gravel, weathering, and small frost wedges at the top
of 22 of the 40 units led Chambers (1971, 1984) to conclude that
glacial Lake Missoula drained at least 22 times to below 959 m above
sea level (asl), which is 306e321 m below the highest elevation
reached by the lake. Waitt (1980, 1985) ascribed rhythmically
1
In this paper, we define “unit” as an ensemble of lithofacies recording a single
lake phase from filling through maximum stage to drainage.
bedded slackwater flood sediments in southern Washington to these
drainings and thus inferred that the drainings were catastrophic.
Baker and Bunker (1985), however, argued that the slackwater beds
are not proof that each draining was catastrophic. Levish (1997)
studied exposures of rhythmically bedded sediments similar to
those at Ninemile in Mission Valley, Montana (Fig. 1). He inferred 66
sedimentation units, each of which is recorded by approximately
30e60 varves that thin up-section and, collectively, represent
3240e3610 years of deposition in glacial Lake Missoula. He found no
evidence of unconformities between the units and concluded that
there was no evidence for repeated drainage of glacial Lake Missoula.
Most recently, Smith (2006) described large gravel flood bars that
underlie rhythmic silt and clay beds in several places along the lower
Clark Fork River (Fig. 1). He attributed these features to catastrophic
draining of glacial Lake Missoula and the finer beds at the Ninemile
site to shallower, later stages of the lake when drainings were less
energetic (Smith, 2006; Alho et al., 2010).
This paper describes the stratigraphy and sedimentology of an
important new section of rhythmically stratified sand, silt, and clay
in the heart of the glacial Lake Missoula basin and compares it to the
Ninemile section. The new section, which we refer to as the rail line
section, provides evidence for several tens of substantial drainings
of glacial Lake Missoula. These drainings repeatedly exposed the
lake floor in the Clark Fork River valley at Missoula. Three new
optical ages, two from the rail line site and one from the Ninemile
site, better constrain the timing of glacial Lake Missoula flooding.
2. Section locations, stratigraphy, and sedimentology
The Ninemile section is a 250-m-long exposure along Interstate
Highway 90 near the confluence of Ninemile Creek and Clark Fork
River in western Montana (N 47.0 , W 114.4 ; Fig. 1). The base and
top of the section are, respectively, 936 m and 959 m asl. The base of
the section is approximately 281 m above the base of the former ice
dam, 200 km to the west. The rail line section is a 1.4-km-long,
northwest-trending exposure along the former track of the Chicago
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M.A. Hanson et al. / Quaternary Science Reviews 31 (2012) 67e81
Milwaukee St. Paul and Pacific Railroad west of the city limits of
Missoula and approximately 26 km upstream of the Ninemile
section (N 46.9 , W 114.1 ; Fig.1). The base and top of the section are,
respectively, 963 m and 975 m asl. The highest glacial Lake Missoula
shorelines around Missoula are 1265e1280 m asl, 290e315 m above
the top of the rail line section (Pardee, 1942; Smith, 2006).
69
2.1. Ninemile Creek section
The Ninemile Creek section comprises 34 units of rhythmically
stratified, graded beds of fine sand, silt, and clay (Figs. 2 and 3A)
and, at the base of the section, at least 1 m of coarse to fine sand and
pebble-cobble gravel. The uppermost 3 units were too weathered
Fig. 2. Stratigraphic log of the Ninemile Creek section, Montana.
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to describe in detail. The sediments comprise three lithofacies:
a gravelesand facies, a sandesilt facies, and a silteclay facies. The
sequence of lithofacies in most units, from bottom to top, is gravelesand, sandesilt, and silteclay facies.
2.1.1. Gravelesand facies
This facies comprises thin beds of coarse sand and, in some
cases, granule to medium pebble gravel. It lies at the base of nine of
the 34 units (Figs. 2 and 3B) and ranges from less than 1e20 cm
thick; the thinnest beds are discontinuous. Most gravel clasts are
subrounded to subangular. The gravelesand facies also fills wedgeshaped cracks at the top of the five lowest units (Fig. 3C).
2.1.2. Sandesilt facies
The sandesilt facies is present in each of the 34 rhythmically
stratified units and comprises very fine sand or silt (Figs. 2 and 3D).
Individual occurrences of the facies range from 9e220 cm thick.
Lower contacts are commonly sharp and planar or undulating; 16 of
the contacts are clearly erosional. Some of the lower contacts,
however, are loaded, producing convoluted bedding above or
below the contact (Figs. 2 and 4A).
The sandesilt facies is characterized by planar or undulating
laminae ranging from <1e4 mm thick. About half of the strata of
the sandesilt facies contain ripples (Figs. 2 and 3D); in one case, the
ripples are overturned. Both type-A and type-B ripples are present
Fig. 3. Rhythmically bedded sand, silt, and clay units at the Ninemile section. (A) One unit comprising sandesilt facies (below) and silteclay facies (above). The contact between the
two facies is gradational. Unit is 320 cm thick. Varves thin upward as the silt content decreases. (B) Gravelesand facies at the base of a unit. Sand is w6 cm thick. (C) Wedge-shaped
feature at the top of a unit, infilled with gravel and sand. Feature is w15 cm long. (D) Type-A and type-B climbing ripples in sandesilt facies. (E) Loaded silt laminae and water
escape structures in sandesilt facies.
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(Jopling and Walker, 1968). Most ripples indicate flow to the
northwest; in one instance, flow was to the southeast (Fig. 2). Other
common sedimentary structures in the sandesilt facies are loaded
laminae and beds, diapirs, dish structures, and flame structures
(Figs. 2, 3E, 4A, and B). Thin (w2 mm) clay laminae are present
locally within the facies, commonly near the top and generally
associated with laminated silt.
2.1.3. Silteclay facies
The silteclay facies is also present in each of the 34 rhythmically
stratified units and comprises sets of silt and clay couplets that
resemble varves (Figs. 2 and 4C). The number of couplets per
rhythmically stratified unit ranges from two to 44 and averages 19.
Couplets are 0.3e20 cm thick. Within each unit, couplet thickness
decreases upward from an average of 4.3e0.8 cm (Fig. 3A). The
lower part of each couplet consists of massive silt with few clay
laminae; the upper part is clay, in some cases with thin silt interlaminae (Fig. 4C). Lower contacts of individual couplets are sharp
but the boundary between the silt layer and overlying clay layer in
each couplet is gradational (Fig. 4C). The thickness of the clay layer
changes little upward within a unit; thin couplets have a higher
ratio of clay to silt than thick couplets and the thinnest ones are
approximately 80 percent clay. Thinner couplets also have fewer
71
internal laminations than thicker couplets. Near the top of three of
the units, couplets are separated by anomalously thick (1.5e5 cm)
silt beds (Fig. 4D). The three silt beds have sharp, planar lower and
upper contacts, and include planar and undulating laminae and
ripples.
2.1.4. Correlation with Chambers’ work
Chambers (1971, 1984) counted six more units at the Ninemile
site than we did, although he was only able to describe 32 in detail;
the remaining units were at the top of the exposure and were too
weathered to describe in detail. Based on unit thickness, elevation,
and sedimentological characteristics, such as the presence of the
sand-gravel facies and ripple-drift sequences in the silt, we are able,
with few exceptions, to correlate our units with Chambers’ units.
Just above our 10-m level, Chambers identifies one additional unit.
We are confident, however, in our decision to include his thin unit
(18.5 cm thick) within the unit that underlies it. Above the 17.5-m
level, correlation is more difficult. Chambers’ 22nd unit from the
base is 1.5 m thick, which does not match the thickness of any of our
units. A possible explanation for this discrepancy is that Chambers’
22nd unit represents two of our units, which occur between 17.4 m
and 18.6 m. Chambers also described one more unit than we were
able to near the top of the exposure, below the weathered zone.
Fig. 4. Sedimentary features at the Ninemile section. (A) Loaded lower part of unit with diapiric structure in gravelesand facies and dish structures in the sandesilt facies. (B) Flame
structures within lower part of sandesilt facies. (C) Nine varves with sharp lower contacts, gradational internal contacts, and internal laminations. (D) Event beds at the top of
silteclay facies.
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Finally, above the 17.5-m level, Chambers also has one occurrence
of sand and gravel, whereas we identified six. A possible explanation for the discrepancy is the discontinuous nature of this facies.
Chambers’ total varve count is 716; ours is 583. His counts are
higher for nearly all units that we can confidently correlate. We
offer two explanations for this discrepancy: (1) weathering of the
upper part of the exposure over the past several decades did not
allow us to observe all of the varves in the uppermost units; (2) we
erred by grouping varves into composite varves, or Chambers erred
by subdividing composite varves. The latter explanation is more
likely in the sandesilt facies, where thin clay laminae might be
interpreted to be part of a composite varve or the top of a noncomposite varve.
2.2. Rail line section
We studied the rail line section in detail where the exposed
sediments are thickest (11.8 m; Fig. 5). All but the lowest sediments
are well exposed on the northeast-facing side of the railway cut; the
lowest sediments were documented on the southwest-facing side
of the cut. The exposed sediments include 29 upward-thinning
units of rhythmically bedded fine sand, silt, and clay (Figs. 5 and
6A). A zone of highly disturbed sediments approximately 6 m above
the base of the exposure is thicker than the units directly above and
below it (Fig. 5). Judging from the average thickness of adjacent
units, we estimate that the disturbed zone includes 3 units,
bringing the total count to 31. Units consist of three lithofacies
similar to those at the Ninemile site. The three lithofacies, in the
order in which they generally occur from bottom to top, are the
sand, silt, and silteclay facies. We describe the facies in detail
because some of their characteristics differ from those at the
Ninemile site.
2.2.1. Sand facies
The sand facies, which resembles the gravelesand facies at the
Ninemile site, occurs below 10 units, all in the lower half of the
exposure (Fig. 5). The facies comprises well-sorted, very fine to fine
sand (Fig. 6A and B). The lower contact of each of the 10 beds is
erosional, and most contacts are planar (Fig. 6B). The facies ranges
in thickness from <1 cm to 6 cm; the thinner occurrences are
discontinuous. The sand is normally graded, commonly rippled at
the top, and, in some cases, laminated at the base. Type-A ripples,
commonly containing silt stringers or clay rip-ups that were eroded
from underlying sediments, indicate flow to the northeast (Fig. 5).
The thick sand at the base of the third lowest unit (Fig. 5) is
different in character from the sand layers at the base of the other
9 units. It is 64 cm thick and mostly massive, but with some
laminations and type-A ripple-drift cross-laminae in the upper
9 cm of the bed.
2.2.2. Silt facies
The silt facies is limited to two beds, 14 cm and 35 cm thick, in
the two lowest units (Fig. 5). The two silt beds resemble most of the
beds of the sandesilt facies at the Ninemile site. Lower contacts are
sharp and erosive. The facies consists of thinly laminated, coarse to
fine silt with type-B ripple-drift cross-laminae (Fig. 7A). A few thin
(w2 mm) clay laminae occur within silt beds near the top of the
facies.
2.2.3. Silteclay facies
The dominant facies at the rail line section is silteclay couplets
similar to those at the Ninemile site. In most of the units, the
silteclay facies either drapes the sand facies (Fig. 6B) or unconformably overlies silteclay couplets of the next lower unit (Fig. 6C).
Fig. 5. Stratigraphic log of the rail line section, Montana. See Fig. 2 for legend.
The number of couplets per unit ranges from 15 to 42 and
averages 24. Couplets range in thickness from 0.4e12 cm. Couplet
thickness decreases upward within each unit from an average of
3.4e0.7 cm (Fig. 6A). The lower contact of each couplet is sharp
(Fig. 6D). Each couplet comprises a lower lamina or bed of
massive silt, in some cases with a few thin laminae of clay near
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M.A. Hanson et al. / Quaternary Science Reviews 31 (2012) 67e81
73
Fig. 6. Rhythmically bedded sand, silt, and clay units at the rail line section. Arrows indicate unit contacts. (A) A sedimentary unit comprising sand facies (below) and silteclay facies
(above). Varves thin upward as the silt content decreases. Square shows location of photo B. (B) Sand facies with type-A ripples and an erosional lower contact. Silt and clay varves
are draped over ripples. (C) The basal sediment of the upper unit unconformably drapes the sediments of the unit below. The uppermost 2 cm of each unit were disturbed during
draining of the lake. (D) Silt and clay varves. Note clay laminations within silt beds. Contacts between couplets and internal contacts are sharp and planar. Undulation of laminations
is caused by post-depositional desiccation.
the top, and an upper clay cap (Fig. 6D). Most contacts between
the massive silt and the clay cap are sharp (Fig. 6D); only in a few
cases are they gradational. The clay caps range in thickness from
0.2e0.7 cm and appear to be normally graded. Few of the clay
caps contain silt laminae. As at the Ninemile site, the ratio of the
thickness of the clay cap to that of the associated silt layer
increases upward within units; the uppermost couplets are
approximately 80 percent clay. Thinner couplets also have fewer
internal laminations. The uppermost 15 cm of ten of the units are
brecciated and include rip-ups of silt and clay couplets (Figs. 5
and 6C). These disturbed zones have a sharp and probably
erosive lower contact.
Silteclay couplets in other glacial Lake Missoula sediments
identical to those described here have been interpreted to be varves
by Chambers (1971, 1984), Waitt (1980), Fritz and Smith (1993), and
Levish (1997). We interpret these silteclay couplets, as well as
those as the Ninemile site, to be varves. The silteclay couplets meet
five criteria of varves: (1) they are thin, regular, fine-grained, and
commonly contain no current structures; (2) there is little or no
fining in the silt layer, whereas the clay cap fines upward; (3) the
thickness of the clay cap changes little from couplet to couplet,
unlike the thickness of the silt layer; (4) couplets thin and fine
laterally in a down-lake direction; and (5) couplets are found at all
elevations within the lake basin (Ashley, 1975, 2002; Smith and
Ashley, 1985).
2.2.4. Boundaries between units
The contacts between successive units at the rail line section
are unconformable. Even where the basal sand facies is absent,
there is evidence of erosion of up to seven varves at the top of
each unit. Tilted and broken varves are also present in the upper
few centimetres of many units. These broken varves form small
concave-up features 2e3 cm wide and less than 2 cm deep
(Fig. 6A and B). Eight units are characterised by small,
downward-tapering clastic wedges that do not cross the contact
with the overlying unit (Figs. 5 and 7B). These wedge-shaped
structures are filled with silt and clay derived from the adjacent sediment; in some cases brecciated fragments of varves
occur within the casts (Fig. 7B). The cracks extend down to
26 cm from unit contacts. The tops of the cracks range from
0.5e15 cm wide (Fig. 7B). Varves in the adjacent sediment are
bent downward at the margins of the cracks, and micro-faults in
sediment adjacent to the largest crack parallel the crack walls
(Fig. 7B).
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Fig. 7. Unique characteristics of sediments at the rail line section. (A) Laminated and rippled silt facies with clay laminations. Arrow indicates lower contact. (B) Largest wedgeshaped crack, measuring 15 by 26 cm. Crack is infilled with adjacent varved silt and clay. Varves are downturned toward the crack. Micro-faulting can be seen to the right of
the wedge-shaped crack (squares). A secondary similar feature from the overlying unit crosscuts the main feature (arrows).
3. Optical dating
Optical dating provides an estimate of the time elapsed since
deposition and burial of quartz or feldspar grains (see Lian and
Roberts, 2006 and Wintle, 2008 for detailed descriptions of the
method). For this work, quartz was prepared from two samples of
the sand facies at the rail line section (samples MAH-3 and MAH-5)
and from one sample of the gravelesand facies at the Ninemile
section (sample MAH-8). These facies were chosen for dating based
on our interpretation that they were deposited in shallow fluvial
environments on the exposed floor of the lake between ponding
events and before lake refilling (see Section 4.1).
3.1. Sample collection and preparation
We collected samples from horizons free of gravel indicative of
high-energy deposition and avoided sediments showing evidence
of post-depositional disturbance that could have mixed grains of
different ages. Each sample was collected by inserting an aluminum
cylinder (w40 cm long and 10 cm in diameter) into a cleaned face;
the sampled sediment included both the target sand bed and
bounding silt and clay. The cylinder then was excavated and its ends
were filled with fabric and taped to keep the contents immobile
and to retain water content.
In the laboratory, the outer 0.5e1 cm of the sediment from each
end of the cylinder was removed under controlled lighting conditions. Some of this sediment was dried and used for dosimetry and
the remainder was discarded. The sand was then separated from
the surrounding silt and clay for optical dating. All sand samples
were treated with hydrochloric acid to dissolve carbonates and
then sieved to isolate the 180e250 mm size fraction. Samples then
were treated with concentrated hydrofluoric acid for 40 min to
dissolve feldspar grains and to etch from quartz grains outer
material that had been affected by alpha radiation in the environment. Quartz grains were separated from heavier minerals using
a sodium polytungstate solution with a density of 2.68 g ml1. The
separated quartz grains were mounted on 9.8-mm-diameter
aluminum discs using silicone oil as an adhesive. In principle, each
aliquot should be as small as possible, but in practice aliquot size is
dependent on the sensitivity of the quartz grains. For our samples,
each aliquot contained 50 to 100 grains, therefore it is likely that
the measured luminescence was dominated by that originating
from fewer than 10 grains (Duller et al., 2000).
3.2. Environmental dose rate determination
The radiation dose absorbed by the sample grains after deposition comes from alpha (a), beta (b), and gamma (g) radiation
produced by decay of U, Th, their daughter products, and 40K. The
radiation absorbed by the mineral grains of interest originates in
the surrounding sediment and within the sampled grains themselves; there is also a contribution from cosmic rays. Samples used
for dosimetry were dried, milled, and sent to a commercial laboratory for neutron-activation analysis to determine 40K, U, and Th
contents; results are presented in Table 1. From these data, dose
rates (in Gy ka1) caused by g and b radiation were calculated using
standard formulae (Aitken, 1985; Berger, 1988; Lian et al., 1995) and
the dose rate conversion factors of Adamiec and Aitken (1998;
Table 2). In each case, we assumed that the U and Th decay chains
have been in secular equilibrium since deposition of the sediment.
This assumption generally has been found to be valid for welldrained mineral sediments (Prescott and Hutton, 1995; Olley
et al., 1996). Pore water in the sediment matrix attenuates or
absorbs radiation differently from mineral matter, and therefore it
has to be accounted for in the dose rate calculations. When the
sediments were overlain by a lake, it is likely they were saturated
with water. The sediments, however, have spent the vast majority
of the time since deposition as well-drained subaerial deposits,
thus we used the as-collected water contents with an appropriate
uncertainty to account for any reasonable variance (Table 1). For
samples near the ground surface, cosmic-ray radiation can
contribute significantly to the environmental dose rate, but its
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75
Table 1
Sample depths, K, U, and Th contents, and water content.
Sample
Site
d (m)a
K (%)b
U (mg/g)b
Th (mg/g)b
Dw c
MAH-3
MAH-5
MAH-8
Rail line
Rail line
Ninemile Creek
7.6
8.5
2.9
3.61 0.17
3.46 0.18
3.29 0.16
3.06 0.12
3.27 0.13
3.18 0.13
13.0 0.5
13.4 0.5
13.0 0.5
0.172 0.017
0.172 0.017
0.115 0.012
a
d: sample depth beneath the ground surface.
K, U, and Th contents were determined using neutron-activation analysis of a dried and milled subsample of the bulk material sampled for dating.
c
Dw: water content ¼ water mass/mineral mass. For each sample, the as-collected water content was used in the dose rate calculation, with an uncertainty that is thought to
account for any reasonable variation at 2s.
b
effect diminishes rapidly with burial depth (Prescott and Hutton,
1994). For the samples dated here, cosmic-ray radiation constitutes less than three percent of the total dose rate (Table 2).
3.3. Equivalent dose determination
Luminescence measurements were made using a Risø TL/OSL
DA-20 reader. The separated quartz grains were stimulated using
light-emitting diodes that delivered 45 mW cm2 of blue
(470 20 nm) light to the sample. Ultraviolet emissions (w350 nm)
were detected by an Electron Tubes Ltd. 9235QB photomultiplier
tube placed behind a 7.5-mm-thick Hoya U-340 optical filter.
Laboratory irradiations were applied with a calibrated 90Sr/90Y bsource mounted on the reader that delivered 6.38 0.12 Gy min1
to fine sand quartz grains mounted on an aluminum substrate.
The equivalent dose (De; Table 2) of each sample was determined using the standard single-aliquot regenerative dose (SAR)
protocol of Murray and Wintle (2000) but with the inclusion of an
infrared wash to remove any luminescence resulting from K-feldspar contamination (Olley et al., 2004; Wintle and Murray, 2006;
see Fig. 8 for details), which typically was found to be less than one
percent of that measured from the quartz fraction. A zero-dose
point was included to assess the severity of preheat-induced
thermal transfer and one of the regenerative-dose points was
repeated to assess the efficacy of the test dose to correct for
sensitivity change. Of the 73, 49, and 48 aliquots prepared for
samples MAH-3, MAH-5, and MAH-8, respectively, 58, 40, and 35
aliquots were accepted.
For each sample, De values determined using SAR were plotted
on radial plots (Galbraith et al., 1999; Galbraith, 2010), which
display the distributions of De values about mean values (Fig. 9).
The De values for the samples are distributed approximately evenly
about their respective mean values and thus show no obvious
discrete De populations. Therefore, a representative De used in the
calculation of sample burial age was determined using the central
age model (CAM) of Galbraith et al. (1999). This model determines
the weighted De and its associated standard error and takes into
account over-dispersion of the data, which is the variation between
individual De values above and beyond that associated with
Table 2
Dose rates, equivalent doses, and optical ages.
Sample
_ c ðGy=kaÞa
D
D_ T ðGy=kaÞb
De (Gy)c
Optical age (ka)d
MAH-3
MAH-5
MAH-8
0.067 0.003
0.060 0.003
0.129 0.006
4.38 0.3
4.32 0.3
4.45 0.3
65.0 2.0
54.6 1.4
67.2 1.1
14.8 0.7
12.6 0.6
15.1 0.6
a _
Dc : dose rate due to cosmic rays, calculated using present burial depths and the
relationship of Prescott and Hutton (1994).
b _
DT : total dose rates (that due to cosmic rays plus that due to a, b, and g radiation). Only an internal dose rate applies for a radiation, and an estimate of 0.03 Gy/
ka was used (Olley et al., 2004).
c
De: equivalent dose found using the central age model.
d
Apparent optical ages and their analytical uncertainties at 1s.
Fig. 8. (A) Flow chart showing the modified single-aliquot regenerative dose (SAR)
method used to date the rail line and Ninemile samples. Li is the luminescence
measured after administration of a regenerative dose, Di. In this case the regenerative
doses (i ¼ 1 to 5) were 15.9, 31.8. 63.6, 127.2, and 0 Gy; the 0 Gy data point is used to
check for thermal transfer resulting from the preheat, which was 220 C for 10 s. An
aliquot was rejected if the luminescence measured at the 0 Gy step was more than 5
percent of that of the natural. Ti is the luminescence measured following a test dose of
12.7 Gy and a cut-heat of 160 C. For each aliquot, the 15.9 Gy dose (arbitrary choice)
was repeated to determine the effectiveness of the test dose for correcting for sensitivity change; only aliquots that had a recycling ratio within 10 percent of unity were
accepted. (B) Example of a luminescence decay curve for an aliquot of sample MAH-3,
which is typical of those generated from natural aliquots of the three samples. The
initial measured signal is dominated by the thermally stable fast component, which
decays to about 10 percent of its initial intensity in w2 s. The inset graph shows doseresponse data for this aliquot, determined from the sensitivity corrected luminescence
(Li/Ti) measured over the first 0.4 s of stimulation, minus that recorded during the last
20 s. The data are fitted with a saturating exponential curve, onto which the natural
luminescence is interpolated in order to determine the equivalent dose (De).
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M.A. Hanson et al. / Quaternary Science Reviews 31 (2012) 67e81
analytical uncertainties (Galbraith et al., 1999; Lian and Roberts,
2006; Jacobs and Roberts, 2007). Even in cases where aliquots
consisted only of grains that were exposed to sufficient sunlight
prior to burial, over-dispersion values can be as high as 20 percent
(see references cited in Lian and Roberts, 2006, p. 2459). All of the
samples that we analyzed have over-dispersion values of approximately 20 percent (sample MAH-3) or much less than 20 percent
(samples MAH-5 and -8; Fig. 9). These relatively low values suggest
that the aliquots used to determine De had not been significantly
contaminated with older grains that were insufficiently exposed to
sunlight prior to burial. The results are consistent with our interpretation that the sands were deposited in a shallow, non-turbid
fluvial environment (see Section 4.1).
It is possible, although unlikely, that the size of the aliquots has
masked the presence of younger or older grain populations. When
one cannot be certain that the luminescence signal measured from
the sampled grains was properly reset prior to burial or when the
nature of the distribution of De values suggests the presence of
more than one age population, it is common to apply the minimum
age model (MAM) of Galbraith et al. (1999). MAM calculations were
performed on all three samples, and in each case the MAM yielded
De values that are consistent with those found using the CAM.
These results support our supposition that the sampled quartz
grains were exposed to adequate sunlight before burial.
To test whether or not the SAR protocol used here is appropriate
for our quartz samples, we performed dose recovery experiments
(Roberts et al., 1999; Murray and Wintle, 2003) on 24 aliquots each
of samples MAH-3 and MAH-8. Aliquots were exposed to natural
sunlight for 10 min to remove the natural light-sensitive luminescence signal and then each aliquot was given a laboratory radiation
dose similar to the De determined previously using the CAM. We
then attempted to recover the laboratory dose using the same SAR
protocol that we used to determine natural De values. In both cases,
the recovered dose was at least 99 percent of the applied dose.
3.4. Optical ages
The two samples collected from the rail line section yielded
burial ages, in calendar years, of 14.8 0.7 ka (MAH-3) and
12.6 0.6 ka (MAH-5; Table 2). The samples were taken from
directly below the 5th and 7th units of glacial Lake Missoula sediments, numbered from the base of the section (Fig. 5). The single
sample from the Ninemile section (MAH-8) returned an age of
15.1 0.6 ka (Table 2). It was collected directly below the 1st unit of
glacial Lake Missoula sediments, as numbered from the base of the
exposure (Fig. 2).
4. Discussion
4.1. Sedimentology
The rhythmic units at the rail line site are similar to those at the
Ninemile site, as well as to those in Mission Valley described by
Levish (1997), except for the lack of unconformities between units
in Mission Valley. Thin gravelesand and sand beds separate
rhythmic units in the Ninemile and rail line sections. These are
likely the gravel-filled channels described by Chambers (1971,
1984). Such sediments can be deposited by underflows, but based
Fig. 9. Radial plots showing the distribution of equivalent doses determined from
aliquots of samples (A) MAH-3, (B) MAH-5, and (C) MAH-8. More precise data plot
farther from the origin. In each case, the gray bands, which are generated by the radial
plot software, are centered on the weighted mean equivalent dose (De) values provided
by the central age model (CAM). De values that plot within the gray bands are
consistent with the mean at a 95 percent confidence interval. De values that showed
poor recycling ratios or unacceptably high thermal transfer (recuperation) were not
included in the CAM calculations. The uncertainty in the weighted mean De used in the
age calculation is provided by CAM. Over-dispersion values are 25 4, 16 2, and
9.5 1.3 percent for samples MAH-3, -5, and -8, respectively.
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M.A. Hanson et al. / Quaternary Science Reviews 31 (2012) 67e81
on evidence for exposure of the underlying unit (see below), we
interpret the gravel and sand to be non-channelized fluvial sediments deposited in shallow water on the floor of the lake before
refilling; Chambers (1971) reached a similar conclusion at the
Ninemile site.
Thick, laminated and rippled basal silt beds are present
throughout the Ninemile section but were found only in the two
lowest units at the rail line section. At Ninemile, these beds have
been interpreted by Shaw et al. (1999) as deposits of turbidity
currents generated by periodic jökulhlaups from beneath the
Cordilleran ice sheet in the Rocky Mountain Trench. This explanation does not explain, however, why upwardly thinning varves
above this facies indicate a deepening lake over time after deposition of the sand and silt. Nor does this explanation account for
what we interpret as evidence for subaerial exposure (see below)
below this facies. The climbing ripple-drift sequences indicate
a high influx of suspended sediment and high accumulation rates.
Soft-sediment deformation structures, which are present at Ninemile but absent from the two similar units at the rail line section,
are evidence of rapid deposition causing differential loading and
consequent dewatering of the sediments (Lowe and LoPiccolo,
1974; Rust, 1977). These characteristics, combined with the
erosive lower contact, leads us to conclude that these two thick silt
beds were likely deposited by turbidity currents but at relatively
shallow depths as glacial Lake Missoula began to fill, a conclusion
also reached by Chambers (1971, 1984) at the Ninemile site.
The sand and silt beds at both sites grade up into silteclay varves
that were deposited in deeper water. The varves thin upward
within each unit, evidence of continued deepening of the lake over
time and increasing the distance from sediment sources. The
general decrease in the number of varves per unit up-section at
each site is consistent with a thinning ice dam toward the end of
the Fraser Glaciation. The 1.5e5-cm-thick silt beds that interrupt
the thinning-upward sequence near the top of some units at
Ninemile (Fig. 4D) are deposits of rare, high-energy events, possibly
turbidity currents. They are not separate, thin, lake-ponding units
because they are abruptly and conformably overlain by varves and
because the varves above the silt beds are similar in thickness and
in clay-silt ratios to those directly below them. Deformed beds at
the top of 10 units at the rail line section (Fig. 6C) are likely related
to the unconformity directly above. We exclude the possibility that
these deformed beds could be caused by rapid loading and sedimentation from underflows because these beds have erosive lower
contacts and are not consistently overlain by the silt facies that
could have produced the loading. These beds possibly result from
erosion and rapid re-deposition of saturated glaciolacustrine sediment during draining of the lake.
We cannot conclusively correlate the rail line and Ninemile
sections on the basis of their stratigraphy and sedimentology. Units
at the base of the rail line section, however, are similar to units
throughout the Ninemile section, indicating that at least the lower
part of the rail line section possibly overlaps with the Ninemile
section, which also appears to be older based on our ages. If there is
no overlap between the two sections, then the 31 units at rail line
are missing from the top of the Ninemile section. In this scenario, it
would be difficult to explain how earlier glacial Lake Missoula units
were preserved at the Ninemile site, but later ones were eroded,
given that the later lake draining events probably were less erosive
because the ice sheet was thinning and the lakes were smaller.
Thus, there is likely some overlap between the two sites.
The broken and tilted varves and the concave-up features at the
top of the rail line units probably record desiccation caused by
subaerial exposure. We tentatively interpret the large wedge-like
cracks as ice-wedge casts because: (1) they are infilled with adjacent varved sediment; (2) adjacent varves are bent down toward
77
the cracks implying secondary infill; and (3) micro-faults occur in
adjacent sediment (Harry and Gozdzik, 1988; Murton and French,
1993; Gao, 2005). Chambers (1971, 1984) described similar cracks
at the Ninemile site that he interpreted to be frost wedges filled
with the gravelesand facies. Shaw et al. (2000) did not observe
cracks at Ninemile, but argued that they could be dewatering
structures associated with loading caused by rapid deposition of
the overlying sandesilt units. The cracks at Ninemile, however, are
filled with the gravelesand facies, which nowhere appears
substantial enough to load the underlying varved sediment. Similarly at the rail line site, what we interpret to be ice-wedge casts are
filled with adjacent sediment, not overlying sediment as would be
expected from injection of liquefied sediment into fractures caused
by rapid sedimentation. Unfortunately, due to the nature of the
exposure, we could not examine these features in plan view, which
bears on our interpretation of them as ice-wedge casts. If the
wedges are ice-wedge casts, the lake floor must have been exposed
between units.
The interpretation of the depositional environment of each of
the facies and of the boundaries between units is subject to
different interpretations, but considering all the evidence together,
we believe we have made a credible case for repeated draining and
filling of glacial Lake Missoula at the Ninemile and rail line sites.
Desiccation in the upper few centimetres and possible ice-wedge
casts at the tops of units indicate subaerial exposure of varved
units. Overlying laminated and rippled sand and gravel represent
initial infilling of the lake; laminated and ripple-drift sand and silt
sequences represent increased sedimentation in a shallow but
deepening lake. Lastly, varved sediment with couplets that
decrease in thickness upward indicate that the lake deepened until
the ice dam failed, after which some of the lake sediment became
disturbed and the lake bottom exposed.
Subaerial exposure at the top of every unit at the rail line section
indicates that glacial Lake Missoula drained or partially drained at
least 29 times. Each time, the lake lowered below 960 m asl. The
existence of possible ice-wedge casts in some units suggests that
the lake floor was exposed for at least several years. This length of
exposure is longer than one would expect from simple lake-level
fluctuations and is consistent with the fact that it might take
several years after the lake drained for the lake to refill to the
elevation of this exposure. Twenty-two of the units at the Ninemile
site show evidence of subaerial exposure, including erosional
contacts, frost wedges, or overlying fluvial sand and gravel. This
evidence indicates that the majority of the lake drainings were
significant enough to lower the lake to below 936 m asl. Loaded
contacts between varves and overlying sandesilt beds at three of
the contacts at Ninemile possibly support the argument that glacial
Lake Missoula did not drain completely each time the ice dam
failed.
The hypothesis of repeated draining of glacial Lake Missoula is
consistent with the conclusions drawn by Chambers (1971, 1984)
and Waitt (1980) at the Ninemile site. It is contrary, however, to
the conclusions of Levish (1997) based on his study of similar glacial
Lake Missoula sediments in Mission Valley. He argued that these
sediments show no evidence of repeated subaerial exposure and
that the silt beds at the base of rhythmic beds could be attributed to
episodic surging of the Flathead lobe of the Cordilleran ice sheet,
which increased sedimentation in Mission Valley. Levish did not
describe any features at the contacts between units that would
indicate subaerial exposure, thus we assume that his interpretation
of continuous sedimentation is correct. Levish’s interpretations are
sedimentologically plausible, but there are two other explanations
for rhythmic sedimentation with no evidence of subaerial exposure
that would be more consistent with evidence of repeated glacial
Lake Missoula draining found elsewhere in the basin. First, it is
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M.A. Hanson et al. / Quaternary Science Reviews 31 (2012) 67e81
possible that the Mission Valley sediments were deposited at
a different time than the sediments at the Ninemile and rail line
sites; Levish’s optical ages (19.2e16.0 ka ago) would seem to
support this interpretation. Levish argued on the basis of regional
topography, however, that no more than approximately 100 years
of sediment could be missing from the top of the record, implying
that the exposures represent the later phases of glacial Lake Missoula, as do the sediments at the Ninemile and rail line sites.
Second, Levish’s sites are in the deepest part of the glacial Lake
Missoula basin and thus conceivably remained inundated when
part of the lake drained. Clarke et al. (1984) considered this possibility when modelling glacial Lake Missoula outburst floods and
found that not all models resulted in complete lake drainage. Some
model runs left up to 22 percent of the water in the lake. A water
depth of greater than 220 m, approximately 35 percent the
maximum lake depth, would have inundated most of the exposures
in Mission Valley (Levish, 1997; Smith, 2006). Significant fluctuations in water level at these sites, caused by partial drainage, may
explain the cyclicity of the units and the rippled sand and silt facies.
In contrast, the rail line and Ninemile sites require a minimum lake
depth of about 370 m, or approximately 60 percent of the
maximum lake depth, to remain inundated during a potential
partial draining and thus would more likely be exposed during such
events.
Although there is strong evidence for repeated draining of
glacial Lake Missoula in both the rail line and Ninemile sections, we
cannot conclusively demonstrate that these drainings were
complete or catastrophic. Excellent preservation of sediment at the
two sites suggests that flow out of the basin at these localities did
not significantly erode the lake floor. Hydraulic modelling by Alho
et al. (2010) indicates that the sediments at the rail line section
could be preserved during most lake drainings, even during higherenergy floods. It is likely, however, that the sediments at Ninemile
would only be preserved during smaller, less energetic lake
draining events (Alho et al., 2010), which is consistent with our
optical ages that situate these exposures toward the end of the
existence of glacial Lake Missoula (see Section 4.2).
4.2. Chronology
Radiocarbon ages associated with the advance of the Cordilleran
ice sheet into eastern Washington, with glacial Lake Missoula flood
sediments in Oregon, with low salinity anomalies inferred from
freshwater diatom abundances off the southern Oregon coast that
have been attributed to glacial Lake Missoula flooding, and with
glacial Lake Missoula sediment in the North Pacific all place the
onset of glacial Lake Missoula flooding at or after 23.0 to 19.0 ka
ago2 (Clague et al., 1980; Benito and O’Connor, 2003; Lopes and
Mix, 2009; Gombiner et al., 2010). Based on varve counts and
optical dating of the clay portion of glacial Lake Missoula varves in
Mission Valley, Levish (1997) concluded that the lake existed from
approximately 19.2e16.0 ka ago. Similarly, on the basis of a radiocarbon age and varve counts, Atwater (1986) argued that the 89
glacial Lake Missoula floods he identified occurred over a period of
2000e3000 years between 19.4 and 14.7 ka ago. Mount St. Helens
set S (Sg and So) tephra, dated to between 18.5 and 13.3 ka ago
(Mullineaux et al., 1975, 1978; Crandell et al., 1981; Baker and
Bunker, 1985; Waitt, 1985; Clynne et al., 2008) lies at the top of
one flood bed, at least 11 flood beds from the top of slackwater
flood sediment sequences in southern Washington (Waitt, 1980,
2
All calibrated radiocarbon ages discussed are reported directly from the source
or have been calibrated with OxCal 4.1 using the IntCal09 calibration curve (Bronk
Ramsey, 2009; Reimer et al., 2009).
1985). Using this tephra and paleomagnetic secular variation
records from three slackwater sediment exposures in Washington,
Clague et al. (2003) placed the period of glacial Lake Missoula
flooding between either 18.5 and 13.1 or 21.0 and 16.1 ka ago,
depending on the age of the tephra and the number of years
between floods. Radiocarbon ages on turbidites in the North Pacific
extend flooding to between 13.0 and 12.6 ka ago (Zuffa et al., 2000).
The last flood from glacial Lake Missoula, however, must have
occurred before approximately 13.7e13.4 ka ago (Kuehn et al.,
2009), when Glacier Peak tephras G and B blanketed parts of
Washington and Montana, including Mission Valley (Fryxell, 1965;
Porter, 1978; Mehringer et al., 1984; Foit et al., 1993; Sperazza et al.,
2002; Hendrix et al., 2004; Kuehn et al., 2009; Hofmann and
Hendrix, 2010). Hofmann and Hendrix (2010) recovered a pine
needle from a varved sediment sequence in Flathead Lake in
northern Montana; it yielded an age of 14,150 150 yr BP. They
inferred that this varve sequence, which dates back to at least
14,475 150 yr BP, was deposited in a proglacial lake dammed by
a terminal moraine, the Polson moraine, that was built after glacial
Lake Missoula ceased to exist.
The three new optical ages reported in this paper contribute to
the discussion of the age and duration of glacial Lake Missoula and
its outburst floods. The stratigraphic positions of the optical ages
are shown in Figs. 2 and 5. The ages indicate that the Ninemile
section is slightly older than the rail line section, consistent with
the stratigraphy, which shows that the lower beds at rail line are
similar to most beds at the Ninemile site. The two ages at the rail
line site do not differ statistically when their analytical uncertainties are taken into account at two standard deviations, although
the mean ages are stratigraphically reversed. We see no reason to
favour one age over the other and thus adjust the age of the upper
sample to that of the lower by assuming that w100 years separate
the two dated horizons. With this adjustment, the weighted mean
of the two ages, referenced to the lower sample, is 13.5 0.5 ka
(Fig. 5).
The main potential source of uncertainty in optical dating arises
from the assumption that effective bleaching of the relevant electron traps resets the measured luminescence signal of all the grains
immediately before final burial. Anomalously old ages would arise
if a significant number of electron traps had not been emptied
before deposition and burial. In particular, dating of glaciolacustrine sediments can be problematic because the grains are
commonly transported in turbid water (Berger et al., 1987; Berger,
1988; Lian and Hicock, 2001). Our strategy was to avoid this
problem by dating sand deposited in fluvial environments on the
exposed floor of glacial Lake Missoula between ponding events. Our
supposition was that the sand grains would be well exposed to
sunlight in this environment before being buried. Rendell et al.
(1994) concluded that the optical signal from quartz and feldspar
grains in relatively clear water approximately 10 m deep is thoroughly reset in 3 h. If, however, the water is turbid, as is common in
proglacial environments, the optical signal might not be reset.
Ditfelsen (1992) found that feldspar grains only 75 cm deep in
turbid water are not reset after 20 h of exposure to light. The
problem is exacerbated by the fact that fluvial grains are not always
equally reset e coarser grains are more likely to be reset than finer
ones (Murray et al., 1995; Olley et al., 1998). The sand grains that we
dated, however, were probably deposited in water no deeper than
10 cm. The sediments sampled for dating came from thin sheets,
rather than thick channel deposits, and were more likely deposited
by non-channelized sheet flow on the exposed lake floor than in
channels. The sands are well sorted and thus were probably
deposited in relatively clear water. The radial plots (Fig. 9) do not
display multiple age populations that would indicate a significant
number of grains had not been properly reset, as would be expected
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M.A. Hanson et al. / Quaternary Science Reviews 31 (2012) 67e81
in turbid water, thus reaffirming our interpretation of the depositional environment of these facies. If these sediments had been
deposited by turbidity currents, we would expect a larger spread
(over-dispersion) in the single-aliquot De values.
It is possible that our optical ages underestimate the true burial
ages of the sediments of interest, in spite of the fact that our quartz
samples passed all quality-control tests mentioned above. Reports
of this scenario are rare, but one example was provided recently by
Steffen et al. (2009) who attempted to date quartz collected from
alluvial fan and terrace sediments in Peru using a SAR protocol
similar to the one that we used in this study. Steffen et al. (2009)
concluded that their calculated quartz ages were too low when
compared to independent age information, probably due to the
influence of an unstable component in the measured luminescence
signal. This unwanted component can be seen in their luminescence decay data, and it is especially evident when compared to
that measured from a standard, well-behaved quartz sample
(Steffen et al., 2009, their Fig. 9). It is clear, however, that the
luminescence decay data recorded from our samples (a representative curve is shown in Fig. 8B) closely resembles that of the
standard quartz sample; both are dominated by the thermally
stable fast component over the first w2 s of stimulation. It is
therefore unlikely that our samples suffer from this effect.
Despite these favourable results, additional confidence in our
optical ages could be attained via three supplementary methods:
(1) comparison of our ages with those derived from chemically and
physically similar quartz sediment in the region for which the age
has been determined previously by an independent method or
methods; (2) the analysis of single grains of quartz, and subsequently looking for and eliminating data derived from grains that
show abnormal luminescence behaviour (cf. Jacobs et al., 2006;
Demuro et al., 2008). Additionally, if single-grain ages can be
successfully determined, then the finite mixture model (Roberts
et al., 2000; Jacobs and Roberts, 2007; Arnold and Roberts, 2009)
may be applied to distinguish age populations, if there are any, from
the sample. Single-grain analysis may, however, be impractical due
to the sensitivity of the quartz in the region; and (3) comparison of
our quartz ages with those derived from the feldspar fractions of
the samples (cf. Steffen et al., 2009) after correction for anomalous
fading.
The three optical ages lie within the range of ages that constrain
the time when glacial Lake Missoula could exist (21.4e13.4 ka ago).
They are near the young end of this age range, which is consistent
with the stratigraphy (Smith, 2006). The rail line samples, for
example, are from beneath the 25th- and 27th-to-the-last glacial
Lake Missoula sediment units there.
Three wells drilled on the surface adjacent to the top of the rail
line section indicate that glacial Lake Missoula sediments extend
another 30e40 m below the base of the section (www.mbmggwic.
mtech.edu). The units at rail line systematically increase in thickness downward and this trend probably continues below the base
of the section. Thus a reasonable estimate of the number of units
below the base of the section, based on the average thickness of the
lowest 5 units (0.80 m), is 38e50, for a total of 69e81 units in the
Missoula basin. These numbers are consistent with Levish (1997)
and Atwater (1986) estimates of, respectively, the number of
glacial Lake Missoula units and flood events.
The total number of varves exposed at the rail line section, based
on the average count for each unit, adding a reasonable number of
varves for the disturbed units in the middle of the section, and
assuming that the lower silt facies capped by clay is one varve, is
732. The total number of varves at the Ninemile section is 583.
These are minimum estimates, because some varves probably were
eroded from the top of each unit and because they do not include
the length of time it took the ice dam to re-establish itself and the
79
lake to refill to the elevations of the sections. Ice-wedge casts can
commence growth within a year after a surface is exposed and
expand up to 3 cm per year (Mackay and Burn, 2002), thus indicating that some surfaces were exposed for a few years.
Estimated refilling times for glacial Lake Missoula range from
w20 years to 110 years. Atwater (1986) counted the number of
varves between successive glacial Lake Missoula flood beds in
a sequence of glacial Lake Columbia sediments. The numbers
decrease from 45 to 55 near the base of the sequence to two
between the last two flood beds. Similarly, Waitt (1984, 1985)
counted 20e55 varves between glacial Lake Missoula flood beds
in three exposures in Idaho and Washington. Many of these counts
are minima for the number of years between floods because they
do not include varves eroded by the floodwaters. Waitt (1985)
calculated a refilling time for a maximum glacial Lake Missoula of
less than 65 years. He based this estimate on the assumption that
the rate of discharge into glacial Lake Missoula during the Fraser
Glaciation was twice the present-day rate of discharge into the
basin. He also estimated that it would take 40 years to fill glacial
Lake Missoula to half its maximum volume. Waitt (1985) estimates
of refilling times seem reasonable in light of the glacial Lake
Columbia varve counts of Waitt (1984, 1985) and Atwater (1986).
Levish (1997) counted 20e107 varves per unit in glacial Lake
Missoula sediments in Mission Valley; most units have 30e60
varves. Assuming that the cyclicity of the units in Mission Valley
corresponds to the draining cycle of glacial Lake Missoula, Levish’s
data provide constraints on glacial Lake Missoula refilling times.
Precise stratigraphic correlation of the rail line section and the
upper part of Levish's (1997) Landslide Bend section is not possible,
but based on our previous discussion, we argue that these exposures record the same time interval. In an attempt to improve our
estimate of the age of the rail line sediments, we attempted to
correlate the 31 rail line units and the last 41 Landslide Bend units
based on the number of varves per unit. In this scenario, all units at
the rail line site have fewer varves than corresponding units at
Landslide Bend owing to erosion during lake draining. Some gaps
between units at rail line are required to make this scenario
feasible. The last 41 units at Landslide Bend record 1348 years of
glacial Lake Missoula sedimentation. If the overlap is valid, 616
varves are missing from the rail line site, an average of 20 varves per
unit. It is likely, however, that there are also missing units at the top
of the rail line section because of erosion and weathering. Using
this number and the weighted mean of the two optical ages, the
base and top of the rail line section would date to, respectively,
13.7 0.5 and 12.4 0.5 ka ago. We make a similar provisional
match of the 34 Ninemile units with units 18e51, numbered from
the top, at Landslide Bend, which record 1505 years of glacial Lake
Missoula sedimentation. Assuming this amount of time is recorded
at Ninemile, the top of that section would date to approximately
13.6 0.6 ka ago.
Our chronology accords with the known maximum age for the
advance of the Cordilleran ice sheet into northern Idaho (21.4e19.6
ka; Clague et al., 1980) and with an age that pre-dates glacial Lake
Missoula flooding in Oregon (23.3e22.3 ka; Benito and O’Connor,
2003). It is consistent with radiocarbon ages on glacial Lake Missoula sediment in the Pacific Ocean that indicate flooding occurred
between 19.6 and 12.6 ka ago (Zuffa et al., 2000). Although on the
young side, our chronology is consistent with Smith's (2006)
interpretation that the glacial Lake Missoula rhythmic silt and
clay deposits date to the later stages of the lake. Waitt (1980, 1985)
inferred, and Atwater (1987) attempted, bed-for-bed correlations
between beds at Ninemile and, respectively, the Touchet Beds and
alternating glaciolacustrine and glacial Lake Missoula flood beds in
glacial Lake Columbia. Based on the age of the set S tephra, known
limits of the Cordilleran ice sheet advance and retreat, and the
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M.A. Hanson et al. / Quaternary Science Reviews 31 (2012) 67e81
estimated period between floods, Waitt (1985) postulated that
glacial Lake Missoula existed for 2000e2500 years between w18.8
and 14.2 ka ago, which is fairly consistent with our chronology.
Based on varve counts and constrained by a radiocarbon age in
glacial Lake Columbia sediment, Atwater (1986) proposed that the
Ninemile beds correlate with flood beds in glacial Lake Columbia
that were deposited during the early half of his record (19.4ew17.1
ka), which is older than the inferred ages of our sections. Our
chronology corresponds well with the younger of the intervals
proposed by Clague et al. (2003) based on paleomagnetic secular
variation correlations and different ages of the Mount St. Helens set
S tephra (w18.5e13.1 ka). The inferred ages of the two exposures
are considerably younger than the optical ages of Levish (1997;
19.2e16.0 ka) and Lopes and Mix (2009) chronology for low salinity
anomalies in the Pacific Ocean that are likely due to inputs of water
from glacial Lake Missoula (19.0e17.0 ka). Levish dated glaciolacustrine sediments, and it seems likely that a significant number of
grains in his samples were not sufficiently reset prior to burial,
leading to apparently old ages. Radiocarbon-dated material in the
cores analysed by Lopes and Mix (2009) could be considerably
older than the sediment in which it lies. Our chronology is also not
consistent with Hofmann and Hendrix's (2010) calibrated radiocarbon age of 14,150 150 yr BP from glaciolacustrine sediments in
Flathead Lake. It is possible, if not likely, that the needle they dated
is older than the enclosing sediments. Notably, the age assigned to
the top of the rail line section is younger than the most likely age for
Glacier Peak G and B tephras (13.7e13.4 ka; Kuehn et al., 2009)
before which glacial Lake Missoula had ceased to exist. This
discrepancy may reflect our inability to precisely define the intervals represented by these two exposures because of the amount of
time missing from each unit due to erosion and lake refilling. The
sediment at the rail line site was deposited near the end of the
Fraser Glaciation. As the ice sheet began to retreat, a thinning ice
dam would have produced more frequent floods, decreasing the
flood recurrence interval and thus the number of years per unit
toward the top of the exposures. Unfortunately, there is no way to
accurately model this decreasing recurrence interval.
5. Summary
The rail line section in Missoula, Montana, exposes 29 units of
rhythmically bedded sand, silt, and clay similar to other deposits
that have been attributed to deposition in glacial Lake Missoula.
Each of the units records partial or complete emptying of glacial
Lake Missoula, and the entire sequence spans approximately 1350
years. The Ninemile section exposes 34 units of rhythmically
bedded sand, silt, and clay and spans approximately 1500 years.
Recurrent draining and subaerial exposure of the lake floor at both
sites are indicated by unconformities, desiccation, possible icewedge casts, and fluvial gravel and sand beds. Three new optical
ages limit the time of marine oxygen isotope stage 2 glacial Lake
Missoula flooding. Based on a tentative correlation between our
two sites and Levish's (1997) Landslide Bend site, the Ninemile
section spans the period 15.1 to 13.6 0.6 ka ago, and the rail line
section spans the period 13.7 to 12.4 0.5 ka ago.
Acknowledgements
Our research was supported by the Natural Sciences and Engineering Research Council (NSERC) of Canada through Discovery
grants to Clague and Lian and through an NSERC PGS B Scholarship
to Hanson. M. Al Suwaidi, J. Dykstra, and C. Reiss assisted in the
field, and N. Cox, L. Day, and W. Moon assisted with analyses in the
Luminescence Dating Laboratory at University of the Fraser Valley.
Thoughtful and careful reviews by V. Baker and an anonymous
reviewer greatly improved the paper.
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