(This is a sample cover image for this issue. The actual cover is not yet available at this time.)
This article appeared in a journal published by Elsevier. The attached
copy is furnished to the author for internal non-commercial research
and education use, including for instruction at the authors institution
and sharing with colleagues.
Other uses, including reproduction and distribution, or selling or
licensing copies, or posting to personal, institutional or third party
websites are prohibited.
In most cases authors are permitted to post their version of the
article (e.g. in Word or Tex form) to their personal website or
institutional repository. Authors requiring further information
regarding Elsevier’s archiving and manuscript policies are
encouraged to visit:
http://www.elsevier.com/copyright
Author's personal copy
Quaternary Science Reviews 62 (2013) 114e141
Contents lists available at SciVerse ScienceDirect
Quaternary Science Reviews
journal homepage: www.elsevier.com/locate/quascirev
The Last Glacial Maximum at 44! S documented by a
Lake Ohau, Southern Alps of New Zealand
10
Be moraine chronology at
Aaron E. Putnam a, b, *, Joerg M. Schaefer a, c, George H. Denton b, David J.A. Barrell d, Sean D. Birkel b,
Bjørn G. Andersen e,1, Michael R. Kaplan a, Robert C. Finkel f, Roseanne Schwartz a, Alice M. Doughty g
a
Lamont-Doherty Earth Observatory, 61 Rt. 9W, Palisades, NY 10944, USA
Department of Earth Sciences, Climate Change Institute, University of Maine, Orono, ME 04469, USA
Department of Earth and Environmental Sciences, Columbia University, New York, NY 10027, USA
d
GNS Science, Private Bag 1930, Dunedin 9054, New Zealand
e
Department of Geosciences, University of Oslo, 0316 Oslo, Norway
f
Department of Earth and Planetary Sciences, University of California, Berkeley, CA 95064, USA
g
Antarctic Research Centre, School of Earth Sciences, Victoria University of Wellington, PO Box 600, Wellington, New Zealand
b
c
a r t i c l e i n f o
a b s t r a c t
Article history:
Received 17 March 2012
Received in revised form
24 October 2012
Accepted 29 October 2012
Available online
Determining whether glaciers registered the classic Last Glacial Maximum (LGM; w26,500ew19,000 yrs
ago) coevally between the hemispheres can help to discriminate among hypothesized drivers of ice-age
climate. Here, we present a record of glacier behavior from the Southern Alps of New Zealand during the
‘local LGM’ (LLGM). We used 10Be surface-exposure dating methods and detailed glacial geomorphologic
mapping to produce a robust chronology of well-preserved terminal moraines deposited during the
LLGM near Lake Ohau on central South Island. We then used a glaciological model to estimate a LLGM
glacier snowline and atmospheric temperature from the Ohau glacier record.
Seventy-three 10Be surface-exposure ages place culminations of terminal moraine construction,
and hence completions of glacier advances to positions outboard of present-day Lake Ohau, at
138,600 " 10,600 yrs, 32,520 " 970 yrs ago, 27,400 " 1300 yrs ago, 22,510 " 660 yrs ago, and
18,220 " 500 yrs ago. Recessional moraines document glacier recession into the Lake Ohau trough by
17,690 " 350 yrs ago. Exposure of an ice-molded bedrock bench located inboard of the innermost LLGM
moraines by 17,380 " 510 yrs ago indicates that the ice tongue had receded about 40% of its overall
length by that time. Comparing our chronology with distances of retreat suggests that the Ohau glacier
terminus receded at a mean net rate of about 77 m yr#1 and its surface lowered by 200 m between 17,690
and 17,380 yrs ago. A long-term continuation of ice retreat in the Ohau glacier catchment is implied by
moraine records at the head of Irishman Stream valley, a tributary of the Ohau glacier valley. The
Irishman Stream cirque glacier advanced to produce a set of Lateglacial moraines at 13,000 " 500 yrs ago,
implying that the cirque glacier was less extensive prior to that advance.
We employed a glaciological model, fit to these mapped and dated LLGM moraines, to derive snowline
elevations and temperature parameters from the Ohau glacier record. The modeling experiments
indicate that a snowline lowering of 920 " 50 m and temperature depression of 6.25 " 0.5 ! C below
modern values allows for the Ohau glacier to grow to an equilibrium position within the LGM moraine
belt. Taken together with a glaciological simulation reported from the Irishman Stream valley, snowlines
and temperatures increased by at least w520 m and w3.6 ! C, respectively, between w18,000 and
w13,000 yrs ago.
Climate parameters derived from the Ohau glacier reconstruction are similar to those derived from
glacier records from Patagonia, to air temperature indicators from Antarctica, as well as to sea-surface
temperature and stratification signatures of the Southern Ocean. We think that the best explanation
for the observed southern LLGM is that southern winter duration modulated Southern Ocean sea ice,
which in turn influenced Southern Ocean stratification and made the surface ocean cooler. Orbitally
induced cooling of the Southern Ocean provides an explanation for the LLGM in the Southern Alps having
Keywords:
Last Glacial Maximum
Southern Ocean
Glacier
Geomorphology
10
Be
Surface-exposure dating
Cosmogenic nuclide
Paleoclimatology
Southern Alps
New Zealand
Antarctica
Patagonia
Glaciology
Snowline
Last Glacial termination
* Corresponding author. Lamont-Doherty Earth Observatory, 61 Rt. 9W, Palisades, NY 10944, USA.
E-mail address: aputnam@ldeo.columbia.edu (A.E. Putnam).
1
Deceased.
0277-3791/$ e see front matter ! 2012 Elsevier Ltd. All rights reserved.
http://dx.doi.org/10.1016/j.quascirev.2012.10.034
Author's personal copy
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
115
been coincident with the northern LGM. We argue further that the global effect of North Atlantic stadials
led to disturbance of Southern Ocean stratification, southward shifts of the subtropical front, and retreat
of Southern Alps glaciers. Collapse of Southern Ocean stratification during Heinrich Stadial-1, along with
attendant sea-surface warming, triggered the onset of the Last Glacial termination in the Southern Alps
of New Zealand.
! 2012 Elsevier Ltd. All rights reserved.
1. Introduction
2. Setting
Late Quaternary ice-age cycles have constituted the fundamental pulse beat of Earth’s climate system over the past one
million years, but a complete explanation of these cycles has yet to
emerge. Variations of ice volume, dominated by Northern Hemisphere ice sheets, resemble orbital frequencies in general (Hays
et al., 1976) and the summer insolation intensity signal at 65! N in
particular (Broecker, 1966; Broecker and van Donk, 1970), as was
predicted by Milankovitch (1941). Additional support for
Milankovitch (1941) comes from the close match between the rate
of change of global ice volume and June insolation at 65! N
(Roe, 2006). In particular, Northern Hemisphere ice sheets
most recently attained maximum sizes during a prominent trough
in the boreal summer insolation intensity curve at 26,500e
19,000 yrs ago, known classically as ‘the Last Glacial Maximum’
(LGM; Mix et al., 2001; Clark and Mix, 2002; Clark et al., 2009;
MARGO Project Members, 2009). But there are problems with
regard to the Southern Hemisphere. For example, Hays et al. (1976)
showed that, over the last 300,000 yrs, biological indicators of
Southern Ocean sea-surface temperature (SST) and water-column
stratification fluctuated in phase with northern ice sheets, despite
different signatures of precession-dominated insolation in each
hemisphere. Moreover, Mercer (1984) identified simultaneous
glacier expansion during the classic LGM in both southern and
northern middle-to-high latitudes. Because summer insolation
intensity is out-of-phase during the LGM at those latitudes, the
apparent northesouth synchrony presents a conundrum that
Mercer (1984) called “a fly in the ointment of Milankovitch Theory”
(p. 307). The solution to this problem requires an explanation for
full-glacial climatic conditions in the Southern Hemisphere during
the classic LGM.
To address this issue, we start by documenting the initiation,
duration, magnitude, and termination of the Local Last Glacial
Maximum (LLGM) cold period in the Southern Alps of New Zealand.
We focused on the geomorphology and chronology of an exceptionally well-preserved set of moraines deposited by an ice-age
glacier, referred to here as the Ohau glacier, which occupied the
Lake Ohau valley on the eastern side of the Southern Alps. This
moraine chronology utilizes recent analytical improvements in
10
Be surface-exposure dating (e.g., Schaefer et al., 2009), as well as
a local 10Be production-rate calibration site (Putnam et al., 2010b).
Glaciers in the Southern Alps are regarded as sensitive recorders of
atmospheric temperatures (e.g., Oerlemans, 1997; Anderson and
Mackintosh, 2006; Anderson et al., 2010; Purdie et al., 2011). We
then employ a version of the University of Maine Ice Sheet glaciological model (UMISM) adapted to mountain-glacier reconstruction
to estimate the temperature conditions, relative to modern values,
required to generate ice of LLGM extent in the Ohau valley. Using
the moraine chronology and glacier-climate reconstruction, we
assess potential drivers of Southern Hemisphere ice-age climate. In
this regard, the Southern Alps are well positioned for evaluating
hypothesized mechanisms of global climate change because they
are at the antipode of the North Atlantic region, far distant from
northern ice sheets as well as from sites of North Atlantic Deep
Water formation.
South Island, New Zealand (40e46! S, 164e170! W) is situated
just north of the modern position of the subtropical front (STF) in
the South Pacific Ocean, in the southern westerly wind belt (Fig. 1).
Produced by oblique convergence of the Indo-Australian and Pacific
plates, the Southern Alps form the spine of South Island and reach
altitudes in excess of 2000 m on the main drainage divide (hereafter “Main Divide”). Heavily faulted, uplifted Mesozoic sedimentary rocks constitute the central Southern Alps east of the Main
Divide. Highly indurated quartzofeldspathic greywacke sandstones
and argillite mudstones are among the most common lithologies
(Cox and Barrell, 2007).
The Southern Alps impose a strong föhn effect on the prevailing
westerly winds, inducing 10 m yr#1 or more precipitation (water
equivalent) just northwest of the Main Divide. Precipitation declines
with increasing distance southeast of the Main Divide (Henderson
and Thompson, 1999). More than 3000 inventoried glaciers cover
1158 km2 of the Southern Alps (Chinn et al., 2005). The present-day
snowline is w1500 m above sea level (a.s.l.) just west of the Main
Divide at Franz Josef Glacier, and trends upwards to >2200 m a.s.l.
east of the Main Divide at the Ben Ohau Range (Lamont et al., 1999).
High precipitation totals in ice catchment basins of the central
Southern Alps result in large ice fluxes through outward-flowing
glaciers. Such rapid ice throughput elevates glacier sensitivity to
atmospheric temperature variations, resulting in response times to
climate change of decades or less for many glaciers of the Southern
Alps (e.g., Anderson et al., 2010). Historical records of variations of
glacier lengths show that non-calving glacier termini have fluctuated in-step during the last century on both flanks of the Southern
Alps. Thus, glaciers on opposite sides of the Main Divide register
a common climate signal on decadal-to-centurial timescales (Chinn,
1996; Burrows, 2005; Barrell et al., 2011).
In foreland regions, moraine belts outline former ice tongues
that flowed outward from the Southern Alps during recent glaciations (Suggate, 1990; Barrell, 2011; Barrell et al., 2011), and associated glaciofluvial outwash plains extend coastwards. Northwest
of the Main Divide, post-glacial coastal erosion has truncated some
distal parts of the LLGM moraine belts in some valleys. However,
well-preserved moraine systems rim the LLGM glacial troughs that
contain Lake Tekapo, Lake Pukaki, and Lake Ohau at the western
margin of the Mackenzie basin southeast of the highest sector of
the Southern Alps. The glaciers that occupied these troughs
deposited numerous lateral and terminal moraine ridges during
and following retreat from the LLGM. Moraines formed subsequently during Lateglacial time lie farther toward the heads of
mountain valleys (Kaplan et al., 2010; Putnam et al., 2010a). These
post LLGM landforms provide a basis for assessing glacier and
climate changes during the Last Glacial termination.
3. Geomorphology
3.1. Overview of the Lake Ohau study area
The Ohau glacier tongue was produced by the coalescence of
major tributary glaciers from the Hopkins River and Dobson River
Author's personal copy
116
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
Fig. 1. Right panel: New Zealand with schematic depiction of surrounding physical oceanography (adapted from Carter et al., 1998). Arrows indicate surface-ocean currents, with
purple denoting subtropical currents, red representing flow along the Subtropical Front (STF), and blue illustrating currents associated with the northern boundary of the Antarctic
Circumpolar Current and the Subantarctic Front (SAF). Elevation color scale is inset. Left panel: Wind speed just west of New Zealand plotted versus latitude [derived from ECMWF
ERA-Interim reanalysis data; averaged over the period AD1979 to AD2011 (Dee et al., 2011)]. (For interpretation of the references to color in this figure legend, the reader is referred
to the web version of this article.)
valleys, both of which drain from the Main Divide. The overall
catchment for this system at the LLGM was w900 km2. The Ohau
glacial geomorphological assemblage is shown in Fig. 2 and is
illustrated in Fig. 3. Farthest outboard are remnants of moraines
and outwash older than the LLGM, while the LLGM moraine belts
are more continuous and lie closer to the glacier trough. The
assemblage of moraines and outwash surfaces has been influenced
by the effects of uplift on the northwest side of the tectonically
active Ostler Fault (Fig. 2) (Cox and Barrell, 2007). Outwash plains
spilled through four gaps across the Ostler Fault uplift zone (Amos
et al., 2007; Ghisetti et al., 2007). The Quailburn, Willowbank,
Clearburn, and Ohau River gaps are former glacier meltwater
spillways, with post-glacial drainage from Lake Ohau confined to
the modern Ohau River (Fig. 2). Time-transgressive eastward
meltwater deflection controlled by uplift along the Ostler Fault was
instrumental in preserving outer terminal moraine belts (Amos
et al., 2007, 2010).
We divide glacial landforms near Lake Ohau into two groups on
the basis of morphology and soil development (Cox and Barrell,
2007; Barrell et al., 2011); these groups have been attributed to
the last two major glacialeinterglacial cycles (Amos et al., 2007,
2010; Cox and Barrell, 2007; Barrell et al., 2011). Detailed
geomorphological interpretations are shown in Fig. 2. Our glacial
geomorphology mapping distinguishes topographic ridges of
moraine from all other hummocky or irregular morainal landforms.
In accord with the approach of Sugden and John (1976), we interpret an individual moraine ridge as marking a former position of
the glacier margin, representing either a stillstand or a culmination
of an advance of the ice margin. The material forming the crest of
a moraine ridge was deposited immediately prior to withdrawal of
the ice margin from that ridge.
Lake Ohau stands at 520 m above sea level (a.s.l.), and the Ohau
moraine complex rises to as much as w700 m a.s.l. Pollen records
from a site at similar altitude 20 km northeast of Lake Ohau
(McGlone and Moar, 1998) indicates an early Holocene forest cover
of shrub and small-tree species, but that from mid-Holocene
onwards, natural fire events led to a more open shrubland and
grassland. Nothofagus forest spread into at least some parts of the
area in the late Holocene. During the past w800 years since human
arrival, fire markedly reduced the native forest cover, resulting in
widespread grassland (McGlone and Moar, 1998). Since European
settlement w150 years ago, frequent burning to assist pastoral
farming has produced the modern vegetation of various introduced
grasses, native short tussock (Festuca novae-zelandiae) and scattered patches of native and introduced shrubs. Introduced
mammals, particularly rabbits and sheep, have generated small
areas of soil erosion on the landforms. In recent years, a move away
from burning practices has resulted in numerous wild-spreading
patches of introduced pines and, on bouldery moraine ridges
near Lake Ohau, considerable recolonization by native shrubs and
tussocks. Patches of native forest persist in some tributaries of Lake
Ohau, giving way to subalpine shrub and grassland above the local
natural treeline at w1200 m a.s.l.
3.2. Ohau I landforms
The outermost and hence oldest glacial landform assemblage
(Ohau I) lies south and southeast of Lake Ohau and is mapped in
green and olive in Fig. 2. The Ohau I landscape is gentle and rolling.
Individual moraine ridges are typically w2e5 m high. Channeled
outwash surfaces have a generally subdued form with patches of
loess draped on sheltered surfaces. Boulders are few, and they
commonly have shattered surfaces with spalled material accumulated around their bases.
The southeast sector of this landform assemblage is preserved at
Table Hill, which has been uplifted, and tilted back to the northwest, by movement on the Ostler Fault (Amos et al., 2007; Ghisetti
et al., 2007) (Figs. 2 and 4). At the crest of the Ostler Fault scarp,
which forms the southeastern edge of Table Hill, vertical offset of
the Ohau I landform assemblage is at least w180 m, based on the
Author's personal copy
Fig. 2. Glacial geomorphologic map of Lake Ohau study area. Panel A. Index map of Lake Ohau catchment. Study area location is given on regional New Zealand map (inset). Blue
line delineates catchment. Panel B. Detailed geomorphologic map of the Lake Ohau valley and foreland area. Roman numerals correspond to glacial geomorphologic landform
assemblages described in text. ORG: Ohau River gap. CG: Clearburn gap. WG: Willowbank gap. QG: Quailburn gap. Yellow dots are locations of boulders sampled for 10Be surfaceexposure dating. Geomorphologic legend is inset. Panel C. Map of Irishman Stream valley tributary of the Lake Ohau catchment. Panel D. Detailed geomorphologic map of glacier
landforms at the head of Irishman Stream, adapted from Kaplan et al. (2010). Geomorphologic legend is inset. (For interpretation of the references to color in this figure legend, the
reader is referred to the web version of this article.)
Author's personal copy
118
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
Fig. 3. Panorama of the Ohau terminal moraine complex. View is to the north. Terminal moraine ridges, ice-contact outwash plains, relict braided stream channels, and kettle lakes
occupy the foreground. Lake Ohau is at center distance, with the Hopkins (left) and Dobson (right) river valleys in the center far distance. Ben Ohau Range is in the middle distance
right of center, and Ohau Range is in distance on the left.
height of the scarp and the observation that Ohau IIeIII outwash
lies, and presumably buries Ohau I deposits, on the downthrown
side of the fault. The Table Hill glaciogenic deposits have been
mapped in two regional formations; Balmoral Formation and, for
older deposits on the crest of the fault scarp, Wolds Formation, by
Gair (1967), Cox and Barrell (2007) and Amos et al. (2007).
The southern sector of the Ohau I landform assemblage
comprises subdued moraines and outwash deposits north and west
of Quailburn and Willowbank gaps. The trends of braided stream
channels preserved on the outwash deposits indicate that meltwater flowed through the Willowbank and Quailburn gaps during
construction of the Ohau I moraines.
3.3. Ohau II landforms
The moraines and outwash surfaces constituting the Ohau II
landform assemblage (Figs. 2 and 5), situated inboard of Ohau I
landforms, have a surface form that is distinctively sharper than the
Ohau I landforms. For the most part, the Ohau II landforms are
patchy remnants, mainly of moraine and moraine ridges, preserved
as islands surrounded by extensive Ohau III outwash plains (Figs. 2,
3, and 6). Moraine ridges in the southern portion of the Ohau II
landform complex are broad and discontinuous. They stand 2e5 m
high above neighboring general moraine or outwash surfaces. In
contrast, the outer moraine ridges in the southwestern portion of
the Ohau II landform complex are relatively sharp and continuous,
with ridges typically 5e10 m high. The landforms mostly have
stony surfaces, with rounded gravel common on general moraine.
Scattered on the landforms are greywacke and semischist boulders
(i.e., textural zones I and IIA/IIB, based on the classification used by
Cox and Barrell, 2007).
The Ohau II glaciogenic deposits, along with those of Ohau III,
are correlated with the regional Mt. John Formation (Gair, 1967;
Amos et al., 2007; Cox and Barrell, 2007). The sharp-crested outer
moraine ridges in the southwestern portion of the moraine
complex are associated with aggradational outwash deposits,
implying that these moraines were formed at the culmination of
ice-margin advances (see Fig. 2). Relatively subdued inboard
Fig. 4. Table Hill and Ohau I moraine and outwash complex. View is to east, with Ostler Fault scarp, partly in shadow, delimiting the right (southeast) margin of Table Hill, with
several landslides on the scarp face, and arrays of sinuous slump scars along the scarp crest. Ohau I moraine ridges trend from bottom to top of image on the northwesterly
backtilted Table Hill platform. Fluvially eroded slopes define western and northern limits of Ohau I complex in center foreground and center distance, respectively. Ohau River, the
Ohau and Pukaki hydro-electric canals, and Lake Ruataniwha (artificially dammed) are visible in distance.
Author's personal copy
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
Fig. 5. Photograph of Ohau II (foreground, center-left), Ohau III (middle ground), and
Ohau IV (left-distance) terminal moraine ridges. View is north. Prominent Ohau III icecontact outwash plains with well-preserved relict braided stream channels are visible
at bottom and middle of photograph. Aoraki/Mount Cook is tallest glaciated peak in far
distance.
topography, which includes moraine ridges in the southern sector
of the Ohau II belt, suggests deposition during recession of the
Ohau glacier terminus.
3.4. Ohau III landforms
Situated inboard of, and locally truncating, the Ohau II landforms is the Ohau III geomorphologic assemblage. The Ohau III
glaciogenic deposits are correlated with the regional Mount John
Formation (Gair, 1967; Amos et al., 2007; Cox and Barrell, 2007).
The Ohau III moraine complex rims the southern perimeter of the
Ohau glacier trough, from the eastern foot of the Ohau Range to the
southern base of the Ben Ohau Range (Figs. 2 and 5). There are three
main elements of the Ohau III complex. Sets of prominent, sharpcrested moraine ridges form the outer part of the complex. These
ridges, standing as much as w10 m above the adjacent ground, are
among the tallest and most continuous moraine ridges observed in
the entire study area. A central zone of the Ohau III moraine belt
comprises rather featureless undulating ground moraine topography with numerous shallow lakes or depressions and relatively
few moraine ridges; on the innermost side of Ohau III is a prominent, eroded ice-contact slope. Well-developed outwash plains
119
Fig. 6. Panel A. Oblique aerial photograph with view southwest toward Ohau IV
moraine ridges and associated outwash (foreground), as well as Ohau III moraine
ridges and associated outwash (distance, beyond prominent eroded ice-contact slope
at far margin of Ohau IV outwash plain). Ohau Canal (lower) and Ohau River (center)
drain left from the lake. Panel B. Photograph of Ohau IV outwash plain (middle).
Vantage is west. Prominent Ohau IV terminal moraine ridge is in distance, and sharply
defined eroded-ice-contact slope of innermost Ohau III is in immediate foreground.
Ohau Range defines skyline. Note exceptionally well-preserved relict braided stream
channels on Ohau IV outwash surface.
emanating from the Ohau III moraine belt extend through the
Clearburn gap and the Ohau River gap (Figs. 2 and 5).
The Ostler Fault has produced a w20 m vertical offset of Ohau III
outwash surface at both gaps (Amos et al., 2007, 2010). The fault is
interpreted to be a northwest-dipping listric thrust, on account of
a substantial backtilt of the outwash surfaces, localized within
w2.5 km northwest of the fault scarp. This tilt pattern implies
a rapid northwestward diminution of the tectonic uplift (Davis
et al., 2005; Amos et al., 2007; Ghisetti et al., 2007). As a result,
vertical movement associated with the Ostler Fault is highly
localized along the hanging wall, and uplift at the Ohau IIeV
moraine belts is negligible (e.g., <1 mm yr#1) in regard to
surface-exposure dating calculations.
The middle to inner parts of the Ohau III moraine belt south of
Lake Ohau exhibit a combination of generally featureless undulating topography, scattered discontinuous moraine ridges w1e5 m
high, numerous shallow lakes and an absence of abutting outwash
deposits. This landscape formed during general recession of the
Ohau glacier terminus.
The innermost Ohau III eroded ice-contact slope extends laterally for w13 km, stands w30 m above the Ohau IV outwash surface
(Fig. 6), and approximates the outer margin of the trough into
Author's personal copy
120
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
which the Ohau glacier withdrew, following formation of the Ohau
III moraine/outwash complex. This ice-contact slope could also
represent an ice-margin advance or stillstand.
3.5. Ohau IV landforms
The Ohau IV moraine belt, dominated by prominent, sharpcrested ridges, forms a rim alongside the southern and eastern
shores of the lake (Figs. 2 and 6). The associated outwash plain has
been incised w20e30 m by the modern Ohau River (Fig. 6). Many
sub-rounded to sub-angular greywacke boulders, some exceeding
4 m in height, are embedded in or rest on the moraine ridges, as
well as on the ice-contact heads of Ohau IV meltwater channels and
outwash surfaces. The Ohau IV glaciogenic deposits, along with
those of Ohau V and VI, are correlated with the regional Tekapo
Formation (Gair, 1967; Amos et al., 2007; Cox and Barrell, 2007).
3.6. Ohau V landforms
The innermost features of the Lake Ohau terminal moraine
complex are subdued discontinuous ridges along the southern
shoreline of Lake Ohau that, at their northeastern ends, project into
the Lake Ohau basin (Fig. 2). Two arcuate moraine ridges projecting
northeastward into the lake mark the last positions occupied by the
Ohau glacier before it withdrew into what is now the lake basin.
Geomorphologic relationships indicate that these moraines were
constructed at the edge of newly formed Lake Ohau (Fig. 2).
3.7. Ohau VI landforms
Glaciogenic landforms representing ice retreat from the Ohau V
terminal positions are designated here as Ohau VI. These landforms include well-expressed lateral recessional moraines on the
western side of the Ohau valley near the head of the lake, but
discrete landforms are more difficult to delineate on the eastern
side of the valley, where broad ice-marginal benches span a large
altitudinal range. The upper part of this flight of benches may
reflect recession from older parts of the Ohau glacial landform
sequence (e.g., Ohau III).
Included within the Ohau VI landform assemblage is an icemolded bedrock bench at the junction of the Hopkins and Dobson rivers w24 km north of the Ohau V terminal moraines (Figs. 2
and 7). A patchy veneer of lodgement till and scattered large
greywacke boulders rest on the molded bedrock. This bedrock
bench was exposed as the Ohau glacier receded from the southern
end of the lake toward the head of the Lake Ohau catchment.
3.8. Irishman Stream moraines
At the LLGM, glacier ice from the Irishman Stream tributary
valley merged with the Ohau glacier. Following substantial ice
recession, glacier ice readvanced, producing prominent bouldery
moraines within the cirque at the head of the Irishman Stream
valley (Fig. 8). These moraines have been documented in detail by
Kaplan et al. (2010), and are the youngest glacial geomorphologic
features considered in this study.
3.9. Summary of geomorphological interpretations
We interpret the following sequence of events from the Lake
Ohau glacial geomorphology. The Ohau glacier achieved its greatest
recognized extent prior to the LLGM, as documented by the Ohau I
landform assemblage. Incomplete preservation of the Ohau I
moraine belt, exacerbated by uplift and tilting due to the Ostler
Fault, reduces the certainty with which the ice extent and local
paleogeography can be inferred. After uplift and tilting of the Ohau I
landforms, the Ohau glacier expanded to construct the outboard
Ohau II landforms, followed by ice retreat that produced the
Fig. 7. Oblique aerial photograph of ice-molded landforms near the junction of the Hopkins (background) and Dobson (behind the photographer) rivers. Vantage is west. These
bedrock landforms comprise part of the Ohau VI landform assemblage. Note broad ice-scour troughs and ridges aligned right to left on the crest of the bench, parallel to the Hopkins
valley ice-flow direction, and a dark line in the foreground at the edge of the bench crest, marking a low (w1 m high) Dobson valley recessional lateral moraine. Barrier Range
defines the skyline.
Author's personal copy
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
121
(Fig. 9). Integrated with the geomorphologic mapping study, we
targeted specific moraine crests for sampling based on the
following criteria: (1) crests represent single, distinguishable icemarginal positions; (2) moraines do not appear to have been
reworked by post-depositional processes, including human influence; and (3) landforms are clearly defined and retain pristine
features such as perched erratics and abandoned ice-marginal
outwash fans with preserved braided stream channels (e.g.,
Fig. 6), all indicating little landform change since initial formation.
Where possible, we sampled several boulders on each moraine
ridge. Boulders were selected based on the following
considerations:
(1) Geomorphologic integrity. We assessed the immediate
surroundings of each sampled boulder to ensure there were no
indications of post-depositional disturbance by slope
processes, fluvial reworking, or human activities. Ideal boulders are those rooted in moraine ridge crests, and they occur
away from steep slopes of landforms that may have experienced degradation as well as adjacent hill slopes that may have
shed rock-fall debris onto the moraine area. We generally did
not sample boulders with heights less than 50 cm above
ground level.
(2) Boulder surface integrity. We avoided sampling boulders that
are heavily jointed, fractured, pitted, and/or exhibit evidence of
spalling or granular disintegration. Preference was given to
sampling from the top center of boulders with near-horizontal,
planar, or gently rounded upper surfaces, keeping away from
any steep edges at the boulder sides.
Fig. 8. Irishman Stream moraine set. A.) Overview photograph of glacier landforms at
the head of Irishman Stream valley. Vantage is east. B.) Photograph of outer prominent
bouldery moraine ridge at head of Irishman Stream valley. Vantage is west (depicted
by arrow in Panel A).
inboard Ohau II recessional moraines. A subsequent ice advance,
which in places overran parts of the Ohau II moraine belt, culminated with formation of the near-continuous outer Ohau III
moraine ridges. The inner w3 km of the central portion of Ohau III
moraine belt comprises subdued and discontinuous moraines
formed during recession from the Ohau III maximum. The eroded
ice-contact slope that forms the inner margin of the Ohau III
moraine belt is thought to mark the edge of the glacier trough as it
existed at Ohau III time. Alternatively, this ice-contact slope may
represent an ice-margin advance or stillstand position. The Ohau IV
moraines were formed at the culmination of a subsequent resurgence of the glacier, by which time meltwater outflow had become
confined to the Ohau River gap. Moraine ridges of the inboard Ohau
V moraine belt mark brief pauses in recession from the Ohau IV ice
limit. Subsequent terminus retreat and ice-surface lowering
exposed the bedrock bench near the Hopkins/Dobson junction that
comprises part of the Ohau VI landform assemblage. Recession
continued well into mountain valleys, after which time a glacier at
the head of Irishman Stream tributary valley (and probably other
glaciers in the Ohau catchment headwaters) readvanced to produce
Lateglacial moraines (Kaplan et al., 2010).
4. Methods
4.1. Sample collection
We sampled only boulders on landforms with clear geomorphologic context relating directly to a former glacier margin
Boulder surfaces were sampled using hammer and chisel or the
‘drill-and-blast’ method of Kelly (2003). We used clinometer and
compass to document the surrounding topographic shielding for all
sample locations. We photographed all boulders from several
different perspectives, and we measured the ground-to-sample
height on four sides of every boulder.
4.2.
10
Be extraction and AMS measurements
Samples were processed for 10Be analysis at the LamontDoherty Earth Observatory (LDEO) Cosmogenic Isotope Laboratory using the methods of Schaefer et al. (2009). The LDEO procedure is available online at: http://www.ldeo.columbia.edu/tcn.
Beryllium ratios (10Be/9Be) were measured with the CAMS accelerator at the Lawrence-Livermore National Laboratory. Samples
processed before September 2007 were measured relative to the
KNSTD standard (10Be/9Be ¼ 3.15e#12). Samples processed during
and after September 2007 were measured relative to the 07KNSTD
standard [10Be/9Be ¼ 2.85e#12; based on a revised 10Be half-life of
1.36 " 0.07 Ma (Nishiizumi et al., 2007)]. After measurement,
10
Be/9Be ratios were corrected for residual boron contamination
and 10Be in procedural blanks. All blank corrections were less than
1.5%. All samples measured relative to the KNSTD standard were
subsequently normalized to 07KNSTD by applying a correction
factor of 0.9042 (Nishiizumi et al., 2007).
4.3. Exposure-age calculations
We calculated surface-exposure ages using the sea-level highlatitude (SLHL) 10Be production rates of Putnam et al. (2010b) and
scaling methods of Stone (2000; ‘St’), Desilets et al. (2006; ‘De’),
Dunai (2001; ‘Du’), and Lifton et al. (2005, 2008; ‘Li’), as well as
a version of the Stone (2000) scaling that incorporates a highresolution version of the Lifton et al. (2008) geomagnetic model
(Putnam et al., 2010b), labeled ‘Lm’. Abbreviations follow the
Author's personal copy
122
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
Fig. 9. Examples of boulders sampled for 10Be surface-exposure dating. Panel A. Photograph, with vantage north, of boulder embedded in Ohau VI ground moraine overlying icemolded bedrock. Panel B. Photograph, vantage northwest, of large boulder embedded in top of outermost Ohau IV moraine ridge ice-contact slope. Panel C. Photograph, vantage
west, of boulder resting on top of outermost Ohau III moraine ridge crest. Panel D. Photograph, vantage west-southwest, of large boulder embedded in ice-distal slope of outermost
Ohau III moraine ridge. Panel E. Photograph, vantage north, of boulder resting on subdued Ohau II moraine ridge remnant that is surrounded by Ohau III outwash. Panel F.
Photograph, vantage east, of boulder resting on Ohau II moraine ridge.
nomenclature of Balco et al. (2008). Tectonic and glacioisostatic
uplift rates at the Lake Ohau moraines are assumed to be negligible
(Putnam et al., 2010b) for moraines away from the Ostler Fault (i.e.,
all moraines except for Ohau I). Even for the Ohau I moraine,
a landscape uplift rate of 1 mm yr#1 (the proposed localized uplift
rate on the northwest side of the Ostler Fault scarp adjacent to the
study area; Amos et al., 2010) would decrease ages by less than 1%
(i.e., within the analytical uncertainty of the data, reported below).
Accordingly, we did not apply a correction for the effects of landscape uplift to any of the landforms dated in this study.
Changes in local atmospheric pressures at sea level during the last
ice age (i.e., since boulder deposition) could have resulted in slightly
higher production rates due to equatorward-shifted westerlies and
hence lower atmospheric pressures in southern middle latitudes
(e.g., Staiger et al., 2007). In this regard, a 14C-dated LLGM-age
production-rate test site reported by Putnam et al. (2010b) verified
the suitability of an early-Holocene production-rate calibration data
set for calculating the age of an w18,000-yr surface, indicating that
temporal shifts in average air pressure, if any, have not had any
discernible influence on 10Be production since the end of the LLGM.
This finding is consistent with the modeling results of Staiger et al.
(2007) for the atmosphere over New Zealand. For these reasons we
did not apply a correction for changes in atmospheric pressure.
An erosion rate was not incorporated into age calculations due to
the resistant nature of the greywacke boulder surfaces (Birkeland,
1982; Schaefer et al., 2006, 2009). Furthermore, as no erosion
correction was applied to the production-rate calibration data set of
Putnam et al. (2010b), any common effects of erosion on this rock
type have already been integrated into the SLHL 10Be production
rate. We followed Schaefer et al. (2006, 2009) and Putnam et al.
(2010b) and did not apply a snow-cover correction for the
following reasons: (1) snowfalls are infrequent at elevations below
w1500 m a.s.l. in the Southern Alps, (2) at lower elevations, winter
snow rarely exceeds depths of w1 m and tends to melt away within
days to weeks, and (3) salient boulder tops and moraine ridge crests
are generally windswept and free of snow throughout the winter.
Author's personal copy
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
123
4.4. Glaciological modeling
We constrained a glaciological model to mapped and dated
moraines (Fig. 2) in order to estimate likely LLGM climatic parameters and average equilibrium-line altitude (hereafter, snowline) for
the Ohau glacier. This aspect of work was carried out using the
University of Maine Ice Sheet Model (UMISM), a 2D finite-element
mass and momentum ice dynamics solver with embedded components for calculating isostasy, thermodynamics, sliding, and surface
mass balance (Fastook and Prentice, 1994; Fastook et al., 2008) (see
Appendix). The force balance in UMISM follows the Shallow Ice
Approximation (SIA) (vertically integrated momentum combined
with continuity) where the dominant stress is internal shear, and
where longitudinal stresses are neglected. Although models of this
sort are most commonly used for simulating ice sheets, they are also
applicable in mountain settings for large trunk glaciers flowing over
gently sloping beds (Le Meur and Vincent, 2003; Le Meur et al., 2004,
2007; Leysinger Vieli and Gudmundsson, 2004; Schäfer et al., 2008).
Other SIA modeling studies over domains comparable in scale to the
Ohau valley include the Wasatch (Laabs et al., 2006) and Uinta
(Refsnider et al., 2008) mountains, Utah, USA, and the Wind River
Range, Wyoming, USA (Birkel et al., 2012).
Primary input to UMISM was a 500-m resolution bedrock
elevation map (resampled from 90 m SRTM data) arranged in
a quadrilateral grid of nodal points over the Ohau LLGM ice
catchment. Initial climate was defined by monthly temperature and
precipitation values for New Zealand extracted from the WorldClim
data set (Hijmans et al., 2005). WorldClim is a 1 km gridded
AD1950e2000 climatology derived from the interpolation of
a global array of weather station data using thin plate smoothing
splines. This baseline modern climate input was adjusted in the
LLGM simulations by shifting the temperature seasonal cycle
according to prescribed anomalies (e.g., DT ¼ #5 ! C implies 5 ! C
cooling), and by scaling precipitation (e.g., P ¼ 75% implies 25%
reduction). In order to account for evolving ice topography, the
monthly temperature fields were modified using an atmospheric
lapse rate of #5 ! C km#1, based on modern observations (Norton,
1985). However, we recognize that this is just a working scenario,
as the value may have varied through time. In a sensitivity test of
the LLGM Ohau glacier, we found that increasing the lapse rate
to #6 ! C km#1 required only a þ0.2 ! C DT compensation relative to
the #5 ! C km#1 baseline simulation in order to achieve the
moraine-defined glacier footprint.
The climate signal is transferred to the ice-flow solver through
a calculation of net annual surface mass balance, or the difference
between snow/ice accumulation and ablation. This difference is
obtained using a typical degree-day method (e.g., Braithwaite and
Olesen, 1989). Although degree-day mass-balance schemes lack
the complexity of energy-balance models (which consider spatial
and temporal variations in albedo, insolation, and cloudiness), they
are widely used in glacier modeling studies for their ease of implementation and overall good fit to observational data (Braithwaite,
2011). In our modeling, annual snow accumulation was determined by summing the precipitation amounts for all months when
the temperature, Tmonth, is &0 ! C. Ablation totals were generated by
first summing melting degrees (md) (mdmonth ¼ Tmonth ' number of
days in month) for all months when Tmonth is >0 ! C. Accumulated
snow was then reduced at a prescribed snowmelt rate (mm per md).
If melting degrees remained once all winter snow had been ablated,
ice was in turn melted at an ice-melt rate. We used melt factors for
snow and ice of 4.6 and 7.2 mm md#1, respectively, determined
empirically for New Zealand (Anderson et al., 2006). Fig. 10 shows
how choice of melt rates can influence snowline.
Ice sheet models such as UMISM have changeable parameters
that to some extent control the behavior of ice flow. In order to
Fig. 10. Mass balance versus elevation plots for the Ohau glacier model domain for
three different snow/ice melt rate (SMR/IMR) regimes with base climate at
DT ¼ #6.2 ! C, P ¼ 100%, and atmospheric lapse rate ¼ #5 ! C km#1. A) SMR ¼ 4 IMR ¼ 7,
B) SMR ¼ 4.6 IMR ¼ 7.2, and C) SMR ¼ 5 IMR ¼ 8 (units mm md#1). Snowline, denoted
by blue lines, ranges 1474e1553 m, which affords an uncertainty of "50 m relative to
the 1525 m value derived from New Zealand-specific degree-day factors. Further tests
indicate that a 50 m rise or fall in snowline can be compensated by changing DT
w0.25 ! C. (For interpretation of the references to color in this figure legend, the reader
is referred to the web version of this article.)
Author's personal copy
124
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
produce meaningful paleo-glacier reconstructions, it is therefore
necessary to tune the model so that output is consistent with
existing observational data. The chief ice-flow parameter tuned for
this study was the sliding law constant, B (see Appendix). The value
of B can be adjusted either to lessen or increase the amount of basal
sliding due to melting and refreezing over rough beds. Low values
afford relatively thin, fast-flowing ice, whereas high values afford
the opposite. Another parameter, E, the ice-flow enhancement
factor, can be tuned in order to change ice hardness to account for
impurities. However, we left E at 1 (i.e., no flow enhancement)
because similar net functionality is achieved by varying the sliding
law constant.
Our experimental approach included more than a dozen sensitivity tests wherein DT was adjusted systematically for three
scenarios of ice properties, defined by B ¼ 0.02, 0.03, and 0.04
(Fig. 10). Success or failure was determined by whether or not the
modeled Ohau glacier expanded to LLGM moraine limits. We also
tested cases for two precipitation scenarios, P ¼ 100% and P ¼ 70%.
The latter value accounts for hypothesized drier climate over South
Island during glaciations (Drost et al., 2007; Whittaker et al., 2011).
Taken together, the results from these experiments provide
a means to estimate probable temperature conditions across the
Ohau glacier at the LLGM.
We note that our model is based on modern topography and
lake-floor bathymetry, and it does not incorporate the effects of
buried glacier trough geometry. We therefore take only the Ohau
II/III simulation as generally indicative of LLGM ice extent and do
not attempt to quantify a climatic value for the slightly shorter
glacier length represented by the Ohau IV/V moraine belts. In
addition, because we do not know whether or not the receding
Ohau glacier was in equilibrium with climate when it exposed the
bedrock bench at the Hopkins/Dobson confluence, we did not use
the glacier model to estimate climatic conditions associated with
that recession, but instead characterized the deglaciation of
the Ohau valley using glacier-length changes inferred from
geomorphology.
5. Results
5.1.
10
Be surface-exposure dating
Results from 10Be sample analyses and procedural blank
Be/9Be ratios are given in Table 1a and b, respectively. Procedural
blanks indicate a contribution of 2100e21,600 contaminating 10Be
atoms (Table 1b), with an average contribution of 9760 " 5307
atoms ("1s). Combined, corrections for contaminating 10Be
(determined using procedural blanks) and boron were generally
w1%. Analytical uncertainties were typically 2e3%. Reported
sample 10Be concentration uncertainties have been propagated
with uncertainties in the measured blank 10Be concentration. All
age calculations appear in Table 2, and have been referenced to
calendar years BP (i.e., before AD1950, hereafter ‘yrs’ and ‘yrs ago’)
by subtracting 56, 57, and 60 years for samples collected in AD2006,
AD2007, and AD2010, respectively. Hereafter we discuss ages
calculated using the ‘Lm’ scaling protocol [incorporating the highresolution geomagnetic model of Lifton et al. (2008), following
Putnam et al. (2010b)] because this method yields results that tend
to agree best with independent 14C constraints in the Southern
Alps. A correction for geomagnetic effects integrated over the past
20,000 years is about 0.5%. Choosing alternative scaling methods or
geomagnetic models to calculate ages yields agreement to within
about 1.5%, and therefore would not affect any of the conclusions
below. Uncertainties of elevations measured using differential
global positioning system (GPS) are <1 m and thus have negligible
effects on age uncertainties.
10
All exposure ages are plotted on maps in Fig. 11a and b. Based on
the observed consistency of ages, we combined surface-exposure
dates for each moraine belt (Fig. 12). We used c2 statistics to test
the assumption that multiple 10Be dates from a single landform
assemblage register coeval exposure to the cosmic-ray flux, representing the time of landform construction (Table 3). The c2 test
analyzes whether dispersion among samples in each given data set
can be explained by analytical uncertainties alone, or whether
geomorphological factors, such as multiple episodes of deposition,
predepositional (i.e., ‘inherited’) accumulations of cosmogenic 10Be,
or post-depositional modification processes, have influenced
a particular data set (Balco and Schaefer, 2006). 10Be surfaceexposure ages determined from Ohau IIIeV moraine belts exhibit
low c2 values and hence high levels of internal consistency,
implying that each belt of moraine landforms was constructed
over a period that was within the uncertainty of the 10Be dating
method as applied here.
Where distributions of individual 10Be ages for a group of
associated landforms are approximately normal, we used Chauvenet’s Criterion to test for statistical outliers at the 95% confidence level (Bevington and Robinson, 1992; Dunai, 2010).
Outliers identified by this method were excluded from further
statistical treatment. With the exception of ages from the Ohau I
and II moraine belts, all other sample sets (Ohau III, IV, V, VI)
have approximately normal distributions (Fig. 11). After
removing outliers identified by the Chauvenet method, we
calculated the arithmetic mean and standard deviation to
provide a conservative measure of the landform age and associated uncertainty. We present all derived landform ages and
uncertainties in Table 4.
Finally, systematic uncertainties attending 10Be production-rate
and scaling protocols must be taken into account when evaluating
the full range of acceptable landform ages. Because we use the local
10
Be production rate of Putnam et al. (2010b), scaling uncertainties
are negligible. Choosing a different 10Be production-rate value
would shift all 10Be ages, and hence the whole moraine chronology,
systematically. To illustrate the effects of production-rate uncertainties on the Ohau moraine chronology, we present in Table 4
‘minimum’ and ‘maximum’ possible landform ages calculated
using the upper (þ2.1%) and lower (#2.1%) error bounds, respectively, of the local 10Be production-rate (Putnam et al., 2010b).
Landform ages are given below, and also presented in Table 4. It is
important to note that landform ages should only be compared for
a specific production-rate scenario. In the following presentation
and discussion of landform assemblages, we use the mean age
(third column of Table 4), based on the median value of the
production-rate uncertainty range.
5.1.1. Ohau I
The three ages obtained from Ohau I moraines at Table Hill are
between w130,000 and 150,000 yrs (Figs. 11 and 12; Table 3),
indicating deposition during Marine Isotope Stage 6. However, the
ages are discordant (e.g., c2experimental ¼ 23.42 which is greater than
c2expected of 5.99 at 95% confidence). One possibility is that the range
is due to differential surface erosion of the boulders, thus the oldest
age of 149,500 " 3100 yrs is closest to the true age of the moraine
belt. However, the youngest sample came from a prominent
moraine ridge that lies several ridges inboard of the ridge with the
two older samples. Thus, another possibility is that these moraines
are of composite age, and that the age spread is correct. We
accommodate these possibilities conservatively, by suggesting
a tentative age of 138,600 " 10,600 yrs, based on the arithmetic
mean and standard deviation of the three dates. This age assignment is preliminary, pending further quantitative information of
the age(s) of the Ohau I moraines.
Author's personal copy
Table 1a
Lake Ohau moraine complex surface-exposure sample details and
CAMS
laboratory
no.
Sample
ID
Be data.
Carrier
added (g)a
10
Be/9Be " 1s
(10#14)b
[10Be] " 1s
(104 atoms g#1)c
Average 9Be
current
(mA)d
AMS stde
3.0021
5.0339
5.0195
0.2007
0.2008
0.2014
19.02 " 0.45
22.34 " 0.44
20.60 " 0.67
84.28 " 2.00
99.14 " 1.95
91.80 " 3.02
15.2 (2)
13.6 (3)
15.0 (3)
07KNSTDB1,B2
07KNSTDB1,B2
07KNSTDB1,B2
0.999
0.999
0.999
0.999
0.999
0.987
0.999
0.999
0.999
0.999
4.0076
4.0054
4.0033
4.0186
3.4861
4.0183
4.0032
4.0156
4.0265
3.0082
0.2031
0.1974
0.2000
0.2023
0.1813
0.1995
0.2023
0.2012
0.2002
0.2012
7.42
7.45
5.86
7.23
6.07
5.44
5.14
6.06
6.35
4.73
"
"
"
"
"
"
"
"
"
"
0.22
0.22
0.19
0.27
0.18
0.22
0.19
0.22
0.22
0.15
24.90
24.31
19.30
24.09
20.94
18.13
17.11
20.04
20.82
20.72
"
"
"
"
"
"
"
"
"
"
0.75
0.71
0.62
0.92
0.63
0.73
0.65
0.73
0.74
0.69
11.6
12.8
16.4
11.7
19.1
7.9
10.0
11.8
13.3
15.2
(3)
(3)
(2)
(4)
(4)
(4)
(4)
(4)
(3)
(3)
KNSTDB3
KNSTDB3
07KNSTDB4
KNSTDB3
07KNSTDB5
07KNSTDB4
07KNSTDB4
07KNSTDB4
07KNSTDB4
07KNSTDB1,B2
2.00
0.96
1.49
1.37
2.59
1.88
0.86
1.70
1.19
1.5
1.58
2.37
1.60
1.65
1.36
1.18
2.01
1.92
0.87
1.48
0.88
1.04
0.91
1.05
1.02
1.79
0.998
0.999
0.992
0.993
0.999
0.999
0.999
0.999
0.999
0.996
0.999
0.999
0.999
0.999
0.999
0.999
0.999
0.998
0.999
0.999
0.989
0.999
0.999
0.998
0.999
0.999
7.0417
7.0285
7.0495
7.0649
7.0605
7.0138
7.3478
7.0161
4.1080
4.0176
4.0301
5.0339
5.0195
5.0214
6.0094
5.0152
5.0092
5.0107
5.0771
5.0755
5.0042
5.013
5.0076
6.0114
6.0044
6.0062
0.1989
0.2035
0.1983
0.1980
0.2035
0.2026
0.2021
0.1987
0.2020
0.2018
0.1930
0.2008
0.2014
0.2008
0.1993
0.1983
0.2012
0.2005
0.2015
0.2002
0.1986
0.2015
0.2002
0.1989
0.2009
0.1988
6.77
8.33
8.65
8.60
8.58
8.95
8.97
9.01
5.05
4.81
5.15
5.85
5.61
4.92
6.54
5.79
5.59
5.67
6.12
6.07
6.03
5.68
5.62
6.65
6.77
6.91
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
0.18
0.22
0.21
0.25
0.33
0.25
0.23
0.28
0.15
0.16
0.20
0.22
0.15
0.16
0.21
0.16
0.19
0.19
0.19
0.19
0.21
0.15
0.16
0.22
0.23
0.22
12.59
15.96
16.05
15.96
16.37
17.09
16.31
16.87
16.41
15.96
16.29
15.33
14.84
12.97
14.26
15.03
14.73
14.88
15.99
15.74
15.73
15.04
14.75
14.47
14.90
15.04
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
0.34
0.43
0.40
0.47
0.64
0.48
0.43
0.53
0.50
0.53
0.63
0.59
0.41
0.43
0.47
0.44
0.53
0.52
0.52
0.50
0.54
0.42
0.44
0.48
0.51
0.48
22.2
17.9
20.9
14.2
14.6
14.4
18.3
16.8
10.2
11.2
10.7
10.8
15.6
17.8
15.8
15.4
15.2
15.6
14.7
15.4
12.9
16.1
15.3
14.8
16.5
15.5
(2)
(2)
(2)
(2)
(4)
(3)
(2)
(3)
(5)
(4)
(5)
(2)
(3)
(2)
(2)
(3)
(2)
(2)
(2)
(2)
(2)
(3)
(3)
(2)
(3)
(2)
KNSTDB6
KNSTDB7
KNSTDB6
KNSTDB7
KNSTDB7
KNSTDB8
KNSTDB8
KNSTDB8
KNSTDB3
KNSTDB3
KNSTDB3
07KNSTDB9
07KNSTDB4
07KNSTDB4
07KNSTDB1,2
07KNSTDB1,2
07KNSTDB1,2
07KNSTDB1,2
07KNSTDB9
07KNSTDB9
07KNSTDB9
07KNSTDB10,11
07KNSTDB1,2
07KNSTDB1,2
07KNSTDB1,2
07KNSTDB1,2
2.21
2.07
1.45
1.50
2.07
1.13
2.81
1.55
1.65
1.27
0.999
0.998
0.999
0.998
0.992
0.992
0.999
0.999
0.995
0.999
7.1038
7.0410
7.0320
7.3979
7.0396
7.0048
7.0591
7.0317
7.0789
7.0112
0.2003
0.2011
0.1920
0.1981
0.1981
0.1985
0.2038
0.2013
0.2027
0.2035
6.81
6.77
7.12
7.40
6.62
6.89
6.93
6.42
7.09
6.94
"
"
"
"
"
"
"
"
"
"
0.20
0.21
0.24
0.19
0.18
0.18
0.21
0.19
0.23
0.20
12.65
12.78
12.85
13.05
12.26
12.86
13.19
12.09
13.42
13.29
"
"
"
"
"
"
"
"
"
"
0.38
0.40
0.43
0.34
0.34
0.34
0.41
0.37
0.43
0.39
15.6
15.9
14.8
21.8
21.7
22.3
15.9
15.4
14.3
17.8
(2)
(2)
(4)
(2)
(2)
(2)
(2)
(2)
(2)
(2)
07KNSTDB9
KNSTDB7
KNSTDB7
KNSTDB6
KNSTDB6
KNSTDB6
KNSTDB8
07KNSTDB9
KNSTDB7
KNSTDB8
Longitude
(DD)
Elevation
(m a.s.l.)
Boulder size
(L ' W ' H)
(cm)
Sample
thickness
(cm)
Shielding
correction
#44.297439
#44.297739
#44.298589
169.98659
170.00296
170.00239
605.2
615.4
621.1
240 ' 140 ' 120
250 ' 150 ' 120
400 ' 330 ' 100
2.00
1.19
1.89
0.999
0.999
0.999
#44.35097
#44.3489
#44.34733
#44.34317
#44.29125
#44.31537
#44.31787
#44.3236
#44.32126
#44.341789
169.89591
169.89264
169.90467
169.90149
169.98240
169.95898
169.96605
169.97621
169.94935
169.83328
599
600
586
590
579
569.1
564.5
554.5
570.5
680
230
160
220
320
320
160
370
200
410
200
'
'
'
'
'
'
'
'
'
'
220
100
200
180
190
120
330
140
400
190
80
60
90
60
144
80
120
90
160
60
1.81
0.99
1.65
2.78
1.34
2.01
1.34
1.61
1.45
2.35
#44.25613
#44.25611
#44.24982
#44.25058
#44.31698
#44.32263
#44.32507
#44.32681
#44.33402
#44.33353
#44.33335
#44.30317
#44.3041
#44.30449
#44.30615
#44.30711
#44.30738
#44.33238
#44.33351
#44.33352
#44.33376
#44.33366
#44.33412
#44.322689
#44.323181
#44.323181
169.98217
169.98219
169.97236
169.97101
169.94778
169.94554
169.94361
169.9406
169.88676
169.88824
169.88743
169.95896
169.9588
169.95875
169.95875
169.95823
169.95822
169.87581
169.88021
169.88027
169.88213
169.88232
169.88575
169.87317
169.87543
169.87543
592.1
591.6
582.2
585.9
581.8
574.2
572.5
573.9
605
602
603
582.2
583.4
581.7
577.5
576.9
575.5
620
618
618
613
613
606
614
612
612
190
240
210
100
280
260
220
130
400
350
450
260
400
210
270
350
210
300
300
170
430
190
280
280
230
310
'
'
'
'
'
'
'
'
'
'
'
'
'
'
'
'
'
'
'
'
'
'
'
'
'
'
60 ' 160
170 ' 70
160 ' 100
60 ' 70
150 ' 180
120 ' 140
100 ' 60
100 ' 80
240 ' 170
320 ' 130
320 ' 90
200 ' 200
300 ' 310
200 ' 180
230 ' 180
220 ' 160
110 ' 100
170 ' 80
90 ' 90
110 ' 60
250 ' 130
100 ' 50
210 ' 130
190 ' 70
200 ' 80
225 ' 90
#44.28244
#44.27704
#44.27659
#44.27492
#44.27046
#44.269896
#44.29886
#44.29853
#44.2974
#44.29836
169.94186
169.98217
169.94047
169.93968
169.93488
169.9301
169.90032
169.90072
169.90386
169.90023
557.3
561.4
559.0
561.2
564.1
546.3
583.5
583.6
578.3
575.3
370
210
300
230
280
270
550
390
280
600
'
'
'
'
'
'
'
'
'
'
290 ' 50
160 ' 120
240 ' 120
180 ' 100
90 ' 110
180 ' 110
220 ' 400
260 ' 130
180 ' 210
400 ' 240
'
'
'
'
'
'
'
'
'
'
Quartz
weight (g)
125
(continued on next page)
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
Ohau I moraines
BE24238
OH-07-04
BE24240
OH-07-44
BE24241
OH-07-45
Ohau II moraines
BE23323
OH-06-60a
BE23324
OH-06-61
BE24135
OH-06-62
BE23325
OH-06-63
BE27913
OH-07-01
BE24136
OH-07-16
BE24137
OH-07-18
BE24138
OH-07-19
BE24139
OH-07-21
BE24239
OH-07-26
Ohau III moraines
BE22764
OH-06-04
BE22912
OH-06-05
BE22768
OH-06-16
BE22915
OH-06-18
BE22919
OH-06-44
BE22927
OH-06-45
BE22928
OH-06-46
BE22929
OH-06-50
BE23322
OH-06-59
BE23326
OH-06-72
BE23327
OH-06-73
BE24131
OH-07-08
BE24133
OH-07-10
BE24134
OH-07-11
BE24232
OH-07-13
BE24233
OH-07-14
BE24234
OH-07-15
BE24235
OH-07-31
BE24128
OH-07-33
BE24129
OH-07-34
BE24130
OH-07-36
BE24226
OH-07-37
BE24236
OH-07-38
BE24229
OH-07-39
BE24230
OH-07-40
BE24231
OH-07-41
Ohau IV moraines
BE24125
OH-06-08
BE22913
OH-06-10
BE22914
OH-06-11
BE22765
OH-06-13
BE22766
OH-06-14
BE22767
OH-06-15
BE22924
OH-06-22
BE24126
OH-06-23
BE22916
OH-06-24
BE22926
OH-06-26
10
Latitude
(DD)
Author's personal copy
CAMS
laboratory
no.
Sample
ID
a
[10Be] " 1s
(104 atoms g#1)c
Average 9Be
current
(mA)d
AMS stde
0.21
0.20
0.19
0.21
0.22
0.21
0.20
0.21
0.25
13.55
12.85
11.61
13.41
13.21
12.52
12.57
13.57
12.66
"
"
"
"
"
"
"
"
"
0.41
0.37
0.36
0.39
0.39
0.39
0.38
0.40
0.47
16.5
19.3
15.6
17.5
17.2
12.0
11.6
11.7
13.1
(2)
(2)
(2)
(2)
(2)
(3)
(4)
(4)
(2)
KNSTDB7
KNSTDB6
07KNSTDB9
KNSTDB8
KNSTDB8
KNSTDB8
KNSTDB8
KNSTDB8
KNSTDB8
"
"
"
"
"
"
"
0.20
0.18
0.21
0.15
0.21
0.18
0.16
12.75
12.00
12.48
12.48
13.45
12.52
12.52
"
"
"
"
"
"
"
0.39
0.34
0.39
0.28
0.41
0.33
0.30
17.5
21.5
16.2
20.9
16.4
23.2
12.8
(2)
(2)
(2)
(4)
(2)
(2)
(4)
KNSTDB8
KNSTDB6
KNSTDB7
KNSTDB6
KNSTDB7
KNSTDB6
KNSTDB7
"
"
"
"
"
"
"
"
0.14
0.15
0.16
0.29
0.21
0.16
0.23
0.12
12.50
12.31
12.93
12.95
12.32
12.15
11.94
7.56
"
"
"
"
"
"
"
"
0.22
0.24
0.27
0.46
0.32
0.26
0.27
0.15
21.0
15.8
19.1
17.2
18.8
17.9
16.7
20.0
(4)
(4)
(3)
(4)
(3)
(4)
(4)
(4)
07KNSTDB12
07KNSTDB12
07KNSTDB12
07KNSTDB13
07KNSTDB13
07KNSTDB13
07KNSTDB13
07KNSTDB14,15
360
170
100
220
110
90
150
140
180
3.33
1.92
1.90
2.09
2.61
2.51
2.87
1.68
1.00
0.999
0.998
0.999
0.999
0.999
0.999
0.999
0.999
0.998
7.0471
7.0103
7.0804
7.0090
7.0289
7.1630
7.0079
7.0130
7.0027
0.2037
0.1978
0.1974
0.1937
0.1876
0.2014
0.2017
0.1988
0.2008
7.09
6.92
6.33
7.36
7.50
6.76
6.63
7.25
6.70
"
"
"
"
"
"
"
"
"
290
150
90
220
130
240
140
1.36
1.94
1.35
2.15
1.62
1.78
1.53
0.999
0.998
0.999
0.998
0.995
0.998
0.999
7.0118
7.0158
7.0039
7.0248
7.0228
7.0291
7.0364
0.2037
0.1978
0.1999
0.1962
0.2018
0.1974
0.1993
6.66
6.47
6.62
6.79
7.08
6.77
6.69
0.65
1.30
1.33
1.19
1.86
3.14
1.64
2.19
0.992
0.992
0.992
0.993
0.992
0.992
0.980
0.982
8.0589
7.9369
7.7802
7.7958
8.5528
7.8605
10.7265
10.3647
0.1823
0.1820
0.1825
0.1837
0.1837
0.1826
0.1818
0.1816
8.23
7.99
8.21
8.20
8.55
7.81
10.52
6.37
Elevation
(m a.s.l.)
Boulder size
(L ' W ' H)
(cm)
#44.29664
#44.29227
#44.30176
#44.30282
#44.30214
#44.30577
#44.3067
#44.30368
#44.30026
169.90778
169.93141
169.87808
169.88131
169.89023
169.89184
169.88222
169.86948
169.86005
579.2
546.8
586.2
582.9
581.0
580
580
591
598
400
320
310
410
280
130
310
290
470
'
'
'
'
'
'
'
'
'
250
280
280
110
110
120
130
150
230
'
'
'
'
'
'
'
'
'
#44.29694
#44.29279
#44.29291
#44.28507
#44.28508
#44.28456
#44.30005
169.90352
169.90362
169.90602
169.87905
169.87993
169.87991
169.88623
575.7
554.8
550.1
540.0
539.2
539.0
569.9
420
380
170
730
250
680
240
'
'
'
'
'
'
'
230
220
160
260
210
430
240
'
'
'
'
'
'
'
#44.067938
#44.068170
#44.068120
#44.069717
#44.067300
#44.067294
#44.163124
#44.172895
169.866249
169.865613
169.865572
169.866875
169.865716
169.865681
169.813320
169.813869
721.9
718.8
718.2
694.0
722.3
722.3
644.3
622.4
300
310
250
230
205
205
185
115
'
'
'
'
'
'
'
'
260 ' 50
260 ' 80
200 ' 90
120 ' 50
80 ' 40
115 ' 50
170 ' 60
170 ' 40
Carrier
added (g)a
Be/9Be " 1s
(10#14)b
Shielding
correction
Longitude
(DD)
Quartz
weight (g)
10
Sample
thickness
(cm)
Latitude
(DD)
Carrier 9Be concentration is 996 ppm for all samples except for those labeled OH-10, for which the carrier 9Be concentration is 1024 ppm.
Boron-corrected 10Be/9Be. Ratios are not corrected for background 10Be detected in procedural blanks.
Reported [10Be] values have been corrected for background 10Be detected in procedural blanks.
d 9
Beþ3 measured after the accelerator. Reported currents are averaged over all AMS runs for a given sample. The number of AMS runs is given in parentheses.
e
AMS standards to which respective ratios and concentrations are referenced. Reported 10Be/9Be ratios for KNSTD and 07KNSTD are 3.15 ' 10#12 and 2.85 ' 10#12, respectively (Nishiizumi et al., 2007). ‘B’ refers to corresponding procedural blank listed in Table 1b.
b
c
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
BE22917
OH-06-27
BE22770
OH-06-31
BE24127
OH-06-84
BE22930
OH-06-85
BE22931
OH-06-86
BE23328
OH-06-90
BE23329
OH-06-92
BE23330
OH-06-95
BE23331
OH-06-96
Ohau V moraines
BE22925
OH-06-25
BE22769
OH-06-28
BE22918
OH-06-29
BE22771
OH-06-78
BE22920
OH-06-79
BE22772
OH-06-80
BE22921
OH-06-88
Ohau VI moraines
BE29526
OH-10-01
BE29527
OH-10-02
BE29528
OH-10-03
BE29521
OH-10-04
BE29520
OH-10-05
BE29519
OH-10-06
BE29518
OH-10-07
BE29593
OH-10-08
126
Table 1a (continued )
Author's personal copy
127
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
Table 1b
Procedural blank
10
Be data.
CAMS laboratory no.
Sample ID
Carrier added (g)a
10
BE24237
Blank1
17 Aug 07
(B1)
Blank2
17 Aug 07
(B2)
Blank
06 Jan 07
B3)
Blank2
06 July 07
(B4)
Blank1
10 Nov 09
(B5)
Blank
04 Aug 06
(B6)
Blank
11 Sept 06
(B7)
Blank
03 Oct 06
(B8)
Blank1
06 July 07
(B9)
Blank1
06 Aug 07
(B10)
Blank2
06 Aug 07
(B11)
Blank1
20 April 10
(B12)
Blank1
29 April 10
(B13)
Blank1
18 May 10
(B14)
Blank2
20 May 10
(B15)
0.2002
0.04 " 0.02
0.2000
BE24242
BE23332
BE24140
BE27919
BE22763
BE22922
BE22932
BE24132
BE24214
BE24227
BE29537
BE29513
BE29594
BE29596
Be/9Be " 1s (10#14)b
N10Be " 1s (103 atoms)c
Average 9Be current (mA)d
AMS stde
5.1 " 2.9
13.8 (3)
07KNSTD
0.12 " 0.03
16.4 " 3.6
13.8 (2)
07KNSTD
0.1987
0.04 " 0.02
5.5 " 2.1
10.4 (3)
KNSTD
0.1929
0.06 " 0.03
7.2 " 3.9
17.0 (2)
07KNSTD
0.1789
0.02 " 0.02
2.1 " 2.6
16.7 (4)
07KNSTD
0.1981
0.07 " 0.01
10.2 " 1.8
19.2 (3)
KNSTD
0.1988
0.08 " 0.02
6.8 " 2.3
15.5 (2)
KNSTD
0.2029
0.06 " 0.02
8.7 " 2.2
14.9 (3)
KNSTD
0.2012
0.07 " 0.02
9.8 " 2.6
15.0 (2)
07KNSTD
0.2021
0.07 " 0.03
8.9 " 4.4
17.9 (3)
07KNSTD
0.2015
0.05 " 0.02
6.2 " 2.8
16.1 (3)
07KNSTD
0.1824
0.15 " 0.02
18.7 " 2.9
16.7 (3)
07KNSTD
0.1822
0.17 " 0.02
21.6 " 2.9
17.6 (3)
07KNSTD
0.1816
0.08 " 0.02
10.3 " 2.4
14.1 (3)
07KNSTD
0.2004
0.06 " 0.03
8.9 " 4.1
18.0 (3)
07KNSTD
a
Carrier 9Be concentration is 996 ppm for all samples except for B13, B14, and B15, for which the carrier 9Be concentration is 1024 ppm.
Boron-corrected 10Be/9Be.
c
Total 10Be contamination (in atoms) determined from each procedural blank.
d 9
Beþ3 measured after the accelerator. Reported currents are averaged over all AMS runs for a given sample. The number of AMS runs is given in parentheses.
e
AMS standards to which respective ratios and concentrations are referenced. Reported 10Be/9Be ratios for KNSTD and 07KNSTD are 3.15 ' 10#12 and 2.85 ' 10#12,
respectively (Nishiizumi et al., 2007).
b
5.1.2. Ohau II
Ten ages from the Ohau II moraine ridges range from
26,350 " 1000 to 33,860 " 1040 yrs (Fig. 11) and form a bimodal
distribution with a dominant peak (n ¼ 6) and a younger secondary
mode (n ¼ 4) (Fig. 12). The dominant peak is centered at 32,500 yrs
ago. The full distribution (no outliers removed) has an arithmetic
mean of 30,810 " 2450 yrs (1s) and a median of 31,550 yrs. The
distribution yields a c2 ¼ 49.68 which is greater than c2expected of
16.92 at the 95% confidence level (Table 3), indicating that the data
set includes multiple age populations (Balco and Schaefer, 2006).
As boulder heights do not show a robust correlation to surfaceexposure age (R2 < 0.01; see Section 6.1 below), and because
there is no geomorphological evidence of landform degradation,
we have no basis for inferring that post-depositional geomorphological processes have influenced the Ohau II age distribution. Two
ages (OH-07-16 and OH-07-18) forming the young, secondary
mode of the distribution come from remnants of moraines
preserved as small islands in the Ohau III outwash plain, and it is
conceivable that these remnants represent a younger moraineforming event. The remaining two ages (OH-06-62 and OH-0726) are not in morphostratigraphic order and are considered
untenably young. To derive landform-age estimates we follow
Ward and Wilson (1978) and separate the younger mode containing the two morphostratigraphically young ages from older mode.
This procedure returned an age of 32,520 " 970 yrs (arithmetic
mean of older population " 1s) for the outboard landforms and
a tentative age of 27,400 " 1300 yrs ago (arithmetic mean of two
morphostratigraphically consistent samples of the younger
population " 1s) for inboard moraine remnants.
5.1.3. Ohau III
Twenty-six ages from the Ohau III moraine belt exhibit a high
level of internal consistency (Figs. 11 and 12). The ages of two
samples (OH-06-04 and OH-07-11) are identified as outliers by
Chauvenet’s Criterion and are excluded from further analysis. The
resulting outlier-free distribution yields a c2 of 18.95 (Table 3),
Author's personal copy
128
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
Table 2
10
Be surface-exposure ages (in thousands of cal. yrs before AD1950, labeled ‘ka’, "1s) from the Lake Ohau moraines. Samples marked by a single asterisk (*) next to the sample
ID are tentatively attributed to a younger episode of Ohau II moraine formation. Outlying ages are italicized and marked by a double asterisk (**) next to the sample ID. Bold ages
(‘Lm’) are those discussed in text.
Sample ID
Ohau I moraines
OH-07-04
OH-07-44
OH-07-45
Ohau II moraines
OH-06-60a
OH-06-61
OH-06-62**
OH-06-63
OH-07-01
OH-07-16*
OH-07-18*
OH-07-19
OH-07-21
OH-07-26**
Ohau III moraines
OH-06-04**
OH-06-05
OH-06-16
OH-06-18
OH-06-44
OH-06-45
OH-06-46
OH-06-50
OH-06-59
OH-06-72
OH-06-73
OH-07-08
OH-07-10
OH-07-11**
OH-07-13
OH-07-14
OH-07-15
OH-07-31
OH-07-33
OH-07-34
OH-07-36
OH-07-37
OH-07-38
OH-07-39
OH-07-40
OH-07-41
Ohau IV moraines
OH-06-08**
OH-06-10
OH-06-11
OH-06-13
OH-06-14
OH-06-15
OH-06-22
OH-06-23
OH-06-24
OH-06-26
OH-06-27
OH-06-31
OH-06-84
OH-06-85
OH-06-86
OH-06-90
OH-06-92
OH-06-95
OH-06-96
Ohau V moraines
OH-06-25
OH-06-28
OH-06-29
OH-06-78
OH-06-79**
OH-06-80
OH-06-88
Ohau VI moraines
OH-10-01
OH-10-02
St age (ka)
De age (ka)
Du age (ka)
Li age (ka)
Lm age (ka)
131.3 " 3.2
153.2 " 3.1
141.4 " 4.8
130.1 " 3.2
151.6 " 3.1
139.9 " 4.8
130.7 " 3.2
152.3 " 3.1
140.6 " 4.8
129.7 " 3.2
151.1 " 3.1
139.4 " 4.8
128.3 " 3.2
149.5 " 3.1
138.0 " 4.7
34.30
33.26
29.64
33.66
32.60
28.68
26.68
31.62
32.40
29.58
"
"
"
"
"
"
"
"
"
"
1.05
0.98
0.97
1.30
0.94
1.17
1.02
1.17
1.16
0.99
34.35
33.28
29.61
33.70
32.60
28.68
26.71
31.62
32.41
29.45
"
"
"
"
"
"
"
"
"
"
1.05
0.98
0.97
1.30
0.94
1.17
1.02
1.17
1.16
0.99
34.47
33.42
29.77
33.83
32.75
28.82
26.85
31.79
32.57
29.59
"
"
"
"
"
"
"
"
"
"
1.05
0.99
0.97
1.31
0.95
1.18
1.02
1.17
1.17
0.99
34.40
33.33
29.65
33.76
32.66
28.72
26.76
31.68
32.47
29.45
"
"
"
"
"
"
"
"
"
"
1.05
0.99
0.97
1.31
0.94
1.17
1.02
1.17
1.17
0.99
33.86
32.80
29.22
33.21
32.13
28.28
26.35
31.17
31.94
29.11
"
"
"
"
"
"
"
"
"
"
1.04
0.97
0.95
1.28
0.98
1.15
1.00
1.15
1.15
0.98
17.42
21.96
22.48
22.25
22.93
23.98
22.74
23.63
22.31
21.89
22.25
23.71
22.80
19.94
21.98
23.14
22.84
22.26
23.74
23.46
23.71
22.44
22.13
21.60
22.26
22.58
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
0.48
0.60
0.56
0.66
0.91
0.68
0.60
0.75
0.68
0.74
0.86
0.92
0.64
0.69
0.73
0.68
0.82
0.79
0.77
0.75
0.83
0.63
0.67
0.73
0.76
0.73
17.63
22.07
22.58
22.36
23.03
24.06
22.85
23.72
22.41
22.01
22.35
23.79
22.89
20.12
22.10
23.24
22.94
22.35
23.79
23.52
23.77
22.53
22.23
21.72
22.35
22.66
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
0.49
0.61
0.56
0.66
0.91
0.69
0.61
0.75
0.69
0.74
0.87
0.92
0.64
0.69
0.73
0.69
0.83
0.79
0.78
0.75
0.83
0.63
0.67
0.73
0.77
0.74
17.66
22.14
22.66
22.43
23.11
24.16
22.94
23.82
22.49
22.08
22.43
23.89
22.98
20.17
22.18
23.33
23.03
22.42
23.89
23.61
23.86
22.61
22.31
21.78
22.43
22.74
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
0.49
0.61
0.57
0.66
0.92
0.69
0.61
0.75
0.69
0.75
0.87
0.92
0.65
0.70
0.74
0.69
0.83
0.79
0.78
0.76
0.83
0.63
0.67
0.74
0.77
0.74
17.71
22.13
22.64
22.42
23.08
24.11
22.91
23.78
22.47
22.06
22.41
23.85
22.94
20.19
22.17
23.30
23.00
22.40
23.84
23.57
23.82
22.58
22.29
21.78
22.41
22.72
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
0.49
0.61
0.57
0.66
0.91
0.69
0.61
0.75
0.69
0.74
0.87
0.92
0.64
0.70
0.74
0.69
0.83
0.79
0.78
0.76
0.83
0.63
0.67
0.74
0.77
0.74
17.39
21.78
22.28
22.06
22.72
23.73
22.54
23.40
22.12
21.72
22.06
23.48
22.59
19.84
21.81
22.93
22.63
22.06
23.49
23.22
23.47
22.25
21.94
21.44
22.07
22.38
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
0.48
0.60
0.56
0.65
0.90
0.68
0.60
0.74
0.68
0.73
0.86
0.91
0.63
0.68
0.72
0.68
0.82
0.78
0.77
0.74
0.82
0.62
0.66
0.72
0.76
0.73
19.95
18.16
18.20
18.47
17.47
18.49
18.44
18.54
18.78
18.52
19.09
18.45
17.80
18.67
18.48
17.53
17.65
18.72
17.30
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
0.60
0.57
0.61
0.49
0.48
0.50
0.58
0.58
0.61
0.55
0.58
0.53
0.55
0.55
0.55
0.56
0.54
0.55
0.65
20.15
18.38
18.42
18.68
17.70
18.71
18.65
18.74
18.98
18.73
19.29
18.67
18.01
18.87
18.69
17.75
17.86
18.91
17.51
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
0.61
0.58
0.62
0.50
0.49
0.51
0.59
0.58
0.62
0.55
0.59
0.54
0.56
0.56
0.55
0.56
0.55
0.56
0.66
20.20
18.42
18.46
18.72
17.73
18.75
18.69
18.78
19.02
18.77
19.33
18.72
18.05
18.91
18.73
17.78
17.90
18.95
17.54
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
0.61
0.58
0.62
0.50
0.49
0.51
0.59
0.59
0.62
0.56
0.59
0.54
0.56
0.56
0.55
0.56
0.55
0.56
0.66
20.23
18.46
18.50
18.77
17.78
18.80
18.73
18.82
19.06
18.81
19.37
18.76
18.09
18.96
18.77
17.83
17.94
18.99
17.59
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
0.61
0.58
0.62
0.50
0.50
0.51
0.59
0.59
0.62
0.56
0.59
0.54
0.56
0.56
0.56
0.56
0.55
0.56
0.66
19.86
18.12
18.15
18.41
17.44
18.44
18.39
18.48
18.72
18.47
19.02
18.40
17.76
18.61
18.43
17.50
17.61
18.65
17.27
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
"
0.60
0.57
0.61
0.49
0.49
0.50
0.58
0.58
0.61
0.55
0.58
0.53
0.55
0.55
0.55
0.55
0.54
0.55
0.65
17.77
17.11
17.77
18.05
19.47
18.08
17.55
"
"
"
"
"
"
"
0.54
0.48
0.56
0.41
0.60
0.48
0.43
17.99
17.34
18.00
18.28
19.69
18.31
17.77
"
"
"
"
"
"
"
0.55
0.49
0.57
0.42
0.60
0.49
0.44
18.02
17.37
18.04
18.32
19.74
18.35
17.81
"
"
"
"
"
"
"
0.55
0.49
0.57
0.42
0.60
0.49
0.44
18.07
17.43
18.09
18.37
19.78
18.40
17.86
"
"
"
"
"
"
"
0.55
0.49
0.57
0.42
0.60
0.49
0.44
17.74
17.09
17.74
18.01
19.40
18.04
17.52
"
"
"
"
"
"
"
0.54
0.48
0.56
0.41
0.59
0.48
0.43
17.17 " 0.30
17.00 " 0.34
17.30 " 0.31
17.14 " 0.34
17.30 " 0.31
17.14 " 0.34
17.36 " 0.31
17.21 " 0.34
17.11 " 0.30
16.95 " 0.34
Author's personal copy
129
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
Table 2 (continued )
Sample ID
St age (ka)
De age (ka)
Du age (ka)
Li age (ka)
Lm age (ka)
OH-10-03
OH-10-04
OH-10-05
OH-10-06
OH-10-07
OH-10-08**
17.90
18.25
17.05
16.96
17.77
11.44
18.03
18.38
17.19
17.10
17.94
11.64
18.03
18.39
17.18
17.09
17.96
11.66
18.09
18.45
17.25
17.16
18.02
11.71
17.83
18.17
17.00
16.91
17.72
11.49
"
"
"
"
"
"
0.37
0.66
0.44
0.37
0.40
0.23
"
"
"
"
"
"
0.37
0.67
0.45
0.37
0.41
0.24
which is less than a c2 of 35.17 expected for a normal population at
the 95% confidence level. Because the sample population is normally distributed, and because all measures of central tendency
show close agreement, we use the arithmetic mean " 1s landform
age of 22,510 " 660 yrs for the prominent outer Ohau III moraines
(Figs. 11 and 12). Because only the outer parts of this moraine belt
are bouldery, we do not know whether this age is also representative of the middle to inner, presumed recessional, parts of the
moraine belt.
5.1.4. Ohau IV
Nineteen ages from the outer ridges of the Ohau IV moraine belt
(Fig. 11) are tightly distributed and form an approximately normal
population (Fig. 12). The distribution yields a c2 of 20.48, which is
"
"
"
"
"
"
0.37
0.67
0.44
0.37
0.41
0.24
"
"
"
"
"
"
0.38
0.67
0.45
0.38
0.41
0.24
"
"
"
"
"
"
0.37
0.66
0.44
0.37
0.40
0.23
less than the expected value of 28.87 for a normal distribution of
this size (Table 3). All measures of central tendency agree (Table 3).
We identified one age (OH-06-08) as an outlier using Chauvenet’s
Criterion. After the outlier is removed, we use the arithmetic
mean " 1s to assign a landform age of 18,220 " 500 yrs.
5.1.5. Ohau V
Seven ages from the Ohau V moraine ridges form an approximately normal distribution with one older age (OH-06-79;
19,400 " 590 yrs) forming a shoulder on the older tail of the
summed probability curve (Fig. 12). The c2 determined for this data
set is 10.46, which is less than the c2expected value of 12.59 for
a distribution of this size, indicating that all sample variability can
be explained by analytical uncertainty alone (Table 3). The age of
Fig. 11. Glacial geomorphologic map of Lake Ohau area. Panel A. Southern sector of field area. Panel B. Northern sector of field area. 10Be surface-exposure ages are in white boxes.
Lines indicate sample locations. Ages (black text) are in ka ago. Italics denote ages that are statistical outliers (see text). Red text gives the last four digits of sample ID (Tables 1 and
2). Refer to Fig. 2 for geomorphologic legend and location context within the study area. (For interpretation of the references to color in this figure legend, the reader is referred to
the web version of this article.)
Author's personal copy
130
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
mean " 1s age of 17,380 " 510 yrs for the Ohau VI recessional
morainal landforms.
5.1.7. Chronology of Irishman Stream landforms
Kaplan et al. (2010) reported 37 10Be surface-exposure ages from
the Irishman Stream moraines. All these ages were calculated in the
same fashion as those reported in this study. Kaplan et al. (2010)
assigned an age of 13,000 " 500 yrs ago (arithmetic mean " 1s)
to the outermost Irishman Stream moraine ridge (Figs. 2 and 8).
Younger moraines inboard were formed at 12,200 " 400 yrs ago
and 11,500 " 400 yrs ago.
5.2. Glaciological modeling results
Fig. 11. (continued).
19,400 " 590 yrs that produces the shoulder of the distribution is
identified as an outlier by Chauvenet’s Criterion, and is therefore
excluded from further analysis. After this outlier is removed, we
assign an arithmetic mean " 1s age of 17,690 " 350 yrs for the
Ohau V moraine ridges.
5.1.6. Ohau VI
Eight ages from Ohau VI recessional moraines and erratics on
ice-molded bedrock form an approximately normal distribution,
with one significantly younger age (OH-10-08; 11,490 " 230 yrs)
identified as an outlier by Chauvenet’s Criterion (Fig. 12). The full
age distribution yields a c2experimental of 517.2, which is much
larger than the c2expected value of 14.07, implying that the data set
has been influenced by geomorphologic factors (Table 3). The
boulder from which sample OH-10-08 was collected is near
a talus slope at the base of a cliff face (Fig. 11), and it is possible
that this boulder was emplaced by rock-fall 11,490 yrs ago, rather
than by deposition from the receding glacier. Once this outlier
age is removed the c2experimental for this data set is 8.48, which is
less than the c2expected value of 12.59 for a data set of this size
(Table 3). Based on this refined data set we assign an arithmetic
We conducted a series of 4000-year glaciological simulations to
determine a probable range of DT for the LLGM Ohau glacier under
physically plausible ice-flow regimes. In the first case we used
modern precipitation values. Fig. 13 shows ice area, volume, and
thickness results for optimum solutions R1, R2, and R3 with the
sliding law constant set to 0.02, 0.03, and 0.04, respectively. In each
of these simulations, the Ohau glacier grew to fit lateral and
terminal moraine targets (Fig. 14). We obtained values of
DT ¼ #5.6 ! C in the case of relatively thin, fast flowing ice
(R1), #6.5 ! C for relatively thick, slow-moving ice (R3), and 6.2 ! C
for the middle condition (R2). Modern snowline averaged across
the Ohau catchment in the model lies at 2458 m, and at 1478, 1525,
and 1598 m for simulations R1, R2, and R3, respectively. Respective
snowline depression values are 980, 933, and 860 m, with a midpoint of 920 " 50 m. The latter estimate is consistent with the
875 " 50 m snowline depression estimated by Porter (1975) for the
outer LLGM belt “Mount John” moraines of the nearby Lake Pukaki
glacier trough. In a second set of simulations, we used the boundary
conditions of R2, but reduced precipitation by 30%. From this we
found that ice could meet LLGM lateral and terminal moraine
targets only when DT equaled #6.9 ! C.
Considering the full range of solutions reported above, we take
the mid-point DT value of #6.25 " 0.5 ! C as a reasonable estimate of
atmospheric temperature, relative to modern conditions, in the
Ohau catchment during the LLGM. We note that our estimate of DT
for the Ohau glacier is consistent with the #6 to #6.5 ! C estimate
reported by Golledge et al. (2012) from a recent model of the whole
Southern Alps icefield.
6. Geomorphological and glaciological implications of
chronology
10
Be
6.1. Post-depositional geomorphologic stability of Ohau landforms
Post-depositional moraine degradation can lead to boulder
exhumation, which in turn can complicate interpretations of glacier
history drawn from 10Be surface-exposure ages. Within the Ohau
glacier landform assemblage, 10Be surface-exposure ages determined from boulders from the well-expressed moraine belts Ohau
III, IV, and V show little variability beyond that which can be
explained by analytical uncertainty alone (Table 3; Figs. 12 and 15).
Although sampled boulders range from 50 cm to more than 400 cm
tall, surface-exposure ages show no statistical correlation to
boulder height (Fig. 15). Combined with observations of overall
geomorphological stability of the Ohau landforms, we conclude
that moraine degradation by diffusive processes (cf. Putkonen and
Swanson, 2003; Applegate et al., 2012) has been negligible over at
least the last w30,000 yrs, and possibly since w150,000 yrs ago.
We therefore consider that landform ages derived here and discussed below can be confidently attributed to the timing of fluctuations of the Ohau glacier margin.
Author's personal copy
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
131
Fig. 12. Probability density functions (i.e., ‘camelplots’) for the six Ohau morainal landform assemblages. Center black line is arithmetic mean, while vertical black, red, and green lines
are, respectively, 1s, 2s, 3s uncertainty thresholds. Thin black curves are Gaussian curves representing individual samples. Dotted black line represents summed probability distribution
for data set, including outliers. Thick black curve is summed probability distribution with outliers excluded. Statistics are inset. Note that the younger mode of the Ohau II distribution may
reflect, at least in part, a later episode of moraine formation. (For interpretation of the references to color in this figure legend, the reader is referred to the web version of this article.)
6.2. Chronology of LLGM ice-margin fluctuations at Lake Ohau
The 10Be chronology places the maximum LLGM extent of the
Ohau glacier at 32,520 " 970 yrs ago, and affords inboard mean
moraine dates of 27,400 " 1300 yrs ago, 22,510 " 660 yrs ago,
18,220 " 500 yrs ago, and 17,690 " 350 yrs ago. The w22,500-yr
Ohau III moraine belt served as a target for glaciological
modeling, which returned a snowline depression of 920 " 50 m,
and temperature lowering of 6.25 " 0.5 ! C, relative to modern
values. The prominent geomorphological break separating Ohau III
and IV implies that an interstade preceded construction of the Ohau
IV moraines at 18,220 " 500 yrs ago. Shortly after 18,000 yrs ago,
following construction of the Ohau V moraines, the Ohau glacier
evacuated the Lake Ohau basin. During this retreat, the Ohau glacier
Author's personal copy
132
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
Table 3
Summary statistics for moraine 10Be surface-exposure age distributions. Ages are in cal. yrs before AD1950.
Data set
Ohau I (all)
Ohau II (all)
Ohau II (no
outliers)
Ohau III (all)
Ohau III (no
outliers)
Ohau IV (all)
Ohau IV (no
outliers)
Ohau V (all)
Ohau V (no
outliers)
Ohau VI (all)
Ohau VI (no
outliers)
N
Nremoved
Mean age
(yrs)
"1s
(yrs)
Error-weighted
mean (yrs)
Error-weighted
uncertainty (yrs)
Standard error
of the mean (yrs)
Peak age
(yrs)
Median
age (yrs)
c2experimental
c2expected
c2 test
ðc2experimental
!
< c2expected
(95%)
3
10
6
e
e
4
138,590
30,810
32,520
10,600
2450
970
139,010
30,750
32,540
1990
330
440
7490
770
400
128,720
32,470
32,490
137,980
31,550
32,460
23.42
49.68
3.86
5.99
16.92
11.07
Fail
Fail
Pass
26
24
e
2
22,210
22,510
1280
660
21,960
22,460
140
140
250
140
22,310
22,310
22,260
22,330
134.8
18.95
37.65
35.17
Fail
Pass
19
18
e
1
18,300
18,220
610
500
18,290
18,210
130
130
140
120
18,350
18,350
18,410
18,410
20.48
13.17
28.87
27.59
Pass
Pass
7
6
e
1
17,930
17,690
720
350
17,860
17,700
190
190
270
140
17,770
17,760
17,740
17,740
10.46
2.87
12.59
11.07
Pass
Pass
8
7
e
1
16,650
17,380
2140
510
15,670
17,280
120
140
760
190
17,070
17,070
17,050
17,110
517.2
8.48
14.07
12.59
Fail
Pass
produced the Ohau VI recessional moraines on the valley sides and
exposed the glacially molded bedrock knob near the junction of the
Hopkins and Dobson rivers. If it is assumed that the glacier
terminus was located near the Hopkins/Dobson river confluence
when the bedrock bench became exposed, this would imply
a terminus retreat of w24 km and hence a w40% reduction in
length of the Ohau glacier tongue over the approximately 300 yrs
that had elapsed between construction of the Ohau V and VI
landforms. Over the same time interval, the Ohau glacier surface
lowered at least w200 m to expose the Hopkins/Dobson River
bedrock bench. Therefore, a mean net terminus retreat rate of
w77 m yr#1 and ice-surface lowering rate of 0.7 m yr#1 occurred
between w17,700 and w17,400 yrs ago. Consideration of the
longest possible duration of ice retreat, based on Ohau V and Ohau
VI moraine age error limits, affords minimum net terminus retreat
and ice-surface lowering rates of w20 m yr#1 and 0.18 m yr#1,
respectively.
The Ohau glacier tributaries retracted well back into the high
mountain valleys before a subsequent Lateglacial ice resurgence,
documented by Kaplan et al. (2010) as culminating at
13,000 " 500 yrs ago. Doughty et al. (2012) targeted the Irishman
Stream Lateglacial moraines for glaciological modeling and
determined that a snowline w400 m lower than modern values
and a DT of #2.7 " 0.6 ! C was necessary to grow the modeled
glacier to the outer limit of the Lateglacial moraine belt. Thus, we
consider from these results that the bulk of recession from LLGM
moraine belts toward Lateglacial limits, indicating snowline rise
and atmospheric warming of at least w520 m and w3.6 ! C,
respectively, took place within the first few millennia following
the LLGM in the Lake Ohau drainage system of the Southern Alps
of New Zealand. We note that evidence, if any, for the innermost
limit of ice recession during the first pulse of deglaciation was
masked by glacier readvance in Irishman Stream valley during the
ACR. Thus, because the full distance of glacier recession prior to
Table 4
Mean, maximum, and minimum landform ages based on systematic production-rate uncertainty limits. Production rate limits (PRmax and PRmin) are based on ‘Lm’ scaling and
the production-rate uncertainty of "2.1% (from Putnam et al., 2010b).
Data set
Basis for moraine age
Mean moraine age
(PRmean) (yrs)
Min. moraine age
(PRmax) (yrs)
Max. moraine age
(PRmin) (yrs)
Ohau I (all)
Chose: Mean " 1s;
Basis: Scattered small data set;
considered general minimum age
of moraine belt.
Chose: Mean " 1s after young
mode (n ¼ 4) removed.
Basis: bimodal distribution.
Chose: Mean " 1s after old mode
(n ¼ 6) and two morphostratigraphic
outliers removed.
Basis: bimodal distribution.
Chose: Mean " 1s after two outliers
removed.
Basis: Approx. normal distribution.
Chose: Mean " 1s after one outlier
removed. Basis: Approx. normal
distribution.
Chose: Mean " 1s after one outlier
removed.
Basis: Approx. normal distribution.
Chose: Mean " 1s after one outlier
removed.
Basis: Approx. normal distribution.
138,600 " 10,600
)135,680 " 10,600
&141,500 " 10,820
Ohau II (old mode)
Ohau II (young mode;
outliers removed)
Ohau III (outliers
removed)
Ohau IV (outliers
removed)
Ohau V (outliers
removed)
Ohau VI (outliers
removed)
32,520 " 970
)31,840 " 950
&33,200 " 990
27,400 " 1300
)26,820 " 1270
&27,980 " 1330
22,510 " 660
)22,030 " 650
&22,980 " 680
18,220 " 500
)17,830 " 480
&18,600 " 510
17,690 " 350
)17,320 " 350
&18,060 " 360
17,380 " 510
)17,020 " 500
&17,750 " 520
Author's personal copy
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
133
Augustinus et al., 2011), span the past w30,000 yrs and
complement the Ohau glacier record. These records show
a distinct rise in the percentage of grass pollen, at the expense
of shrub and tree pollen, beginning w29,000 yrs ago. Over the
following w11,000 yrs the percentage of grass pollen varied,
suggesting fluctuations between colder and milder conditions. A
sustained shift in the pollen spectra began between w19,000
and w18,000 yrs ago, with progressive decline in grass pollen in
favor of shrub and tree pollen, which became dominant at these
lowland sites by w17,000 yrs ago. Increasing dominance of
shrub and tree pollen w18,000 to w17,000 yrs ago was
commensurate with the onset of glacier recession from the Ohau
and Pukaki LLGM moraine belts.
7.2. Comparison with records elsewhere in the Southern
Hemisphere
Fig. 13. A) Ice area, B) volume, and C) thickness results for three different ice-flow
regimes defined by the sliding law constant in the Ohau glacier model domain: relatively thin, fast moving ice (R1, B ¼ 0.02, dark gray); moderate ice (R2, B ¼ 0.03, light
gray); and thick, slow-moving ice (R3, B ¼ 0.04, black).
the Lateglacial resurgence can not be discerned from geomorphology, we consider the magnitude of this snowline/temperature
rise to be a minimum value.
7. Discussion
7.1. Comparison to other New Zealand glacier records and
paleovegetation signatures
10
Be dates of the youngest terminal moraines near Lake Ohau
are consistent with those from the youngest LLGM moraines
alongside Lake Pukaki, which lies w20 km northeast of Lake
Ohau. Based on detailed 10Be measurements near Boundary
Stream on the right-lateral moraines of Lake Pukaki, the onset of
ice withdrawal was 18,350 " 390 yrs ago (n ¼ 10, no outliers)
(Putnam et al., 2010b). This age is indistinguishable from that of
the Ohau IV moraine belt (18,220 " 500 yrs) that occurs
alongside Lake Ohau. The 10Be dates therefore suggest that the
Pukaki glacier began a major recession about 18,000 yrs ago in
concert with the Ohau glacier.
Well-documented and well-dated pollen records from
lowland sites at Okarito, alongside the northwest flank of the
Southern Alps on South Island west coast (Vandergoes et al.,
2005; Newnham et al., 2007b), and Auckland, on northern
North Island (Sandiford et al., 2003; Newnham et al., 2007a;
As a test of whether the Ohau glacier record reflects general
glacial conditions in the middle latitudes of the Southern
Hemisphere, we compare the Ohau chronology with chronologies for LLGM glacier advances in mid-latitude South America
(Figs. 16 and 17). In the Chilean Lake District (CLD; 39! e43! S),
14
C ages assayed from plant macrofossils document times at
which vegetated land surfaces were overrun by Andean piedmont glaciers and buried intact by glacial deposits (Denton et al.,
1999b). The Last Glaciation terminal moraine complexes were
formed by at least seven episodes of ice advance between
w36,000 and w18,000 yrs ago (Fig. 15), and together these
events comprise the LLGM of the CLD.
Formation of Ohau LLGM moraine belts match closely in time
with glacier advances into the CLD LLGM moraine belts (Denton
et al., 1999b) (Fig. 17). The age of the Ohau II moraine belt
(w32,500ew27,400 yrs) overlaps with the timing of glacier
advances into the CLD LLGM moraine belt at 34,060 " 325 yrs
ago (Puerto Octay site; n ¼ 5; no outliers), 31,120 " 70 yrs ago
(Bahía Frutillar site; n ¼ 17; 8 outliers excluded), and
27,910 " 210 yrs ago (Canal Tenglo site; n ¼ 6; 2 outliers
excluded). Note that all ages given here for the CLD are from 14C
dates reported in Denton et al. (1999b), converted to calendar
ages using the IntCal09 calibration curve (Reimer et al., 2009). The
Last Glacial advance of the Corcovado ice lobe into the outer
moraine belt on Isla Grande de Chiloe peaked at 18,035 " 215 yrs
ago (Delcahue site; n ¼ 34; 6 outliers excluded) (Denton et al.,
1999b), coinciding with construction of the Ohau IV and V
moraine belts at 18,220 " 500 and 17,690 " 350 yrs ago, respectively. Similar ages of w18,000 yrs were obtained for the Last
Glacial advance into the outer moraine belts of the Ancud, Reloncavi, and Llanquihue glacier lobes of the former Patagonian Ice
Sheet in the CLD (Denton et al., 1999b).
The onset of deglaciation in the CLD (Denton et al., 1999b) began
at approximately the same time as glacier recession at Lakes Ohau
and Pukaki in the Southern Alps. In the CLD region, ice had receded
to within w10 km of modern glacier margins by 14,550 " 620 yrs
ago (Heusser, 1990). Farther south, on the east side of the Southern
Andes at Lago Argentino (w50! S), a readvance of a Southern
Patagonian Icefield outlet glacier culminated at w12,990 " 80 yrs
ago (Strelin et al., 2011). Retreat occurred immediately thereafter,
with ice receding to Holocene limits by w12,000 yrs ago (Kaplan
et al., 2011; Strelin et al., 2011). This Lateglacial signature of Patagonian glacier fluctuations mirrors that of the Irishman Stream
tributary glacier in the Lake Ohau catchment (Kaplan et al., 2010)
and those observed elsewhere in the central Southern Alps
(Putnam et al., 2010a).
Paleoceanographic records (Fig. 16) also exhibit patterns similar
to those defined from the Ohau moraine chronology. Alkenone-
Author's personal copy
134
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
Fig. 14. Visualizations of simulation R2 at 4000 model years: A) ice thickness map, B) mass balance map, and C) shaded polygon surface rendered in Google Earth". Model domain
is fully specified for the Lake Ohau catchment only, and does not yield realistic results for neighboring catchments that are not fully included in the model domain. Elevation
contours on the surface of the Ohau glacier are at 100 m intervals.
and Mg/Ca-inferred SSTs from cores off southern Australia (37! S;
MD03-2611; Calvo et al., 2007) and Chile (41! S; ODP-1233; Lamy
et al., 2004; Kaiser et al., 2005), as well as faunal-based SST
reconstructions from the Indian Ocean (46! S; MD88-770; Labeyrie
et al., 1996; Barrows et al., 2007), southeast Atlantic Ocean (41! S;
TN057-21; Barker et al., 2009), and just west of South Island (42! S;
SO136-GC3; Pelejero et al., 2006; Barrows et al., 2007) all record
the onset of decline towards maximum glacial cold by w30,000 yrs
ago. The STF was equatorward of its present-day position during
the LGM (Bard and Rickaby, 2009; Sikes et al., 2009; De Deckker
et al., 2012). A cold and latitudinally expanded Southern Ocean
persisted, with some variability, until w18,000 yrs ago, after which
time sea-surface warming commenced ubiquitously across the
Southern Ocean (Lamy et al., 2004; Barrows et al., 2007; Calvo
et al., 2007; Lamy et al., 2007; Barker et al., 2009; De Deckker
et al., 2012).
The foraminiferal SST reconstructions from cores TNO57-21 at
41! S in the South Atlantic (Barker et al., 2009) and MD03-2611 at
37! S south of Australia (De Deckker et al., 2012) exhibit sharply
expressed cold periods marking northward excursions of the STF
that coincide with the construction of the Ohau III and IV moraine
belts at w22,500 and w18,200 yrs ago, respectively (Fig. 17). Sea
surface warming and poleward migration of the STF commenced
during the Ohau IVeVI recession phase (just after w18,000 yrs
ago), and which was interrupted by a SST/STF reversal that culminated coevally with construction of the Irishman Stream moraines
at w13,000 yrs ago. A final southern mid-latitude SST warming
phase detected in cores TNO57-21 and MD03-2611 coincided with
recession of the Irishman Stream glacier from Lateglacial to Holocene extents (w13,000e11,700 yrs ago; Kaplan et al., 2010). The
TNO57-21 and MD03-2611 faunal records also capture a prominent
interstade, and southward movement of the STF, from w21,000 to
w19,000 yrs ago that corresponds to the prominent geomorphological break between Ohau III and IV (i.e., after w22,500 and
before w18,000 yrs ago), as well as the ‘Varas Interstade’ at
w21,000 to w19,000 yrs ago in the Chilean Lake District (Mercer,
1972; Denton et al., 1999b). The observed close correspondence
among millennial-scale events detected in SSTs and faunal
assemblages south of Australia (Calvo et al., 2007; De Deckker et al.,
2012), Atlantic SSTs (Barker et al., 2009), the Ohau glacier record
presented here from the southwest Pacific region, and the CLD
glacier record from the southeast Pacific, implies that all regions
monitored a common Southern Hemisphere mid-latitude climate
signal during the LLGM and Last Glacial termination.
Author's personal copy
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
135
Fig. 15. Plots comparing boulder height versus ages normalized to mean moraine age for the six Ohau morainal landform assemblages. Bottom panel shows all data plotted
together. Solid black lines show trend, with respective regression correlation coefficients plotted nearby (R2 values). Dotted black line (and respective R2 value) in Ohau VI panel
gives resulting trend with one exceptionally young outlier age excluded from data set. Note that all data sets show no significant correlation between boulder height and
boulder age.
Author's personal copy
136
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
Fig. 16. Polar projection map of the Southern Hemisphere with locations of records mentioned in text. Arrows indicate ocean currents. Red dots are terrestrial sites discussed in text,
blue dots give key marine sediment core locations, and white dots are locations of pertinent ice cores. Dark blue region illustrates sub-antarctic zone. Schematic physical
oceanography adapted from Brown et al. (2001). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)
Furthermore, some aspects of the Ohau glacier record exhibit
similarities to, as well as important differences from, Antarctic icecore isotope-inferred temperature records (w70e80! S; Petit et al.,
1999; Blunier and Brook, 2001; EPICA Community Members, 2006;
Kawamura et al., 2007; Lemieux-Dudon et al., 2010; see Fig. 16 for
core locations). For example, the EPICA Dome C deuterium signature, placed on the timescale of Lemieux-Dudon et al. (2010),
registered minimum ice-age temperatures over East Antarctica
w34,700ew17,900 yrs ago (Fig. 17), consistent with the duration of
the Ohau LLGM. Antarctic warming commenced at w17,900 yrs ago
(Lemieux-Dudon et al., 2010), at the beginning of the last termination (Denton et al., 2010), in concert with the onset of deglaciation at Lakes Ohau and Pukaki. Resurgence of Southern Alps
glaciers at 13,000 yrs ago fell within the Antarctic Cold Reversal
(ACR) defined from Antarctic ice cores (Kaplan et al., 2010; Putnam
et al., 2010a). Advances into the Ohau LLGM moraine belt also
occurred during a period of persistently low atmospheric CO2
concentrations w30,000ew18,000 yrs ago (Fig. 17; Monnin et al.,
2001; Ahn and Brook, 2008). However, most Antarctic isotopeinferred temperature signatures do not register the strong
millennial-scale variability during the LLGM evident in the Ohau
glacier record, in the CLD glacier record (Denton et al., 1999a,
1999b), and in high-resolution SST reconstructions from southern
middle latitudes (e.g., Barker et al., 2009). One exception is the
d18O-inferred atmospheric temperature signature from the EPICA
Dronning Maud Land core (EPICA Community Members, 2006),
which exhibits warming events that coincide with interstadial
intervals indicated by the Ohau glacier record, the CLD glacier
record, and southern mid-latitude SST patterns (Fig. 17).
Overall, southern mid-latitude SSTs records, as well as highlatitude ice-core signatures of atmospheric temperatures and CO2,
are concordant with the duration of sustained cold atmospheric
temperatures as recorded by glaciers in the Southern Alps and
southern South America (Fig. 17). Millennial-scale climate fluctuations were registered most strongly among mid-latitude records,
with weak or absent counterparts in Antarctic ice-core isotopeinferred temperature records. Overall, most Southern Hemisphere
paleoclimate data suggest that the ocean and atmosphere registered a common climate signal throughout the southern quarter of
the globe during the LGM and beginning of the termination.
Therefore, any proposed drivers of Southern Hemisphere climate
must account for the long duration and the coeval termination of
southern LGM conditions in both the ocean and the atmosphere,
from 37! S to 77! S.
7.3. The LGM in the Southern Hemisphere
In seeking an explanation for Mercer’s (1984) ‘fly in the ointment’ conundrum of an apparently synchronous LGM in the
Northern and Southern Hemispheres, we look towards the
Southern Ocean as a likely causative agent. The Southern Ocean has
a propensity to stratify under cold atmospheric conditions (Gordon,
1991; Sigman et al., 2004), such as may arise from long-duration
southern winters (Huybers and Denton, 2008; Keeling and
Visbeck, 2011). Southern Ocean buoyancy-driven stratification
results, at near-freezing water temperatures, from the density of
seawater being more sensitive to changes in salinity than to
changes in temperature (Sigman et al., 2004, 2010). Wintertime
Author's personal copy
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
137
cooling of surface water initiates sea-ice formation. Brine is rejected
during winter sea-ice formation over the Antarctic continental
shelf (Bouttes et al., 2010, 2011). Subsequent wind-driven export of
sea ice into the open ocean promotes surface-water freshening and
open-ocean stratification (Sigman and Haug, 2003). Thus winter
cooling and sea-ice expansion play essential roles in spreading
stratification into the Southern Ocean. Stratification cools the
surface of the Southern Ocean by inhibiting upwelling of relatively
warm deep waters (Gordon, 1981), enhancing the spread of winter
sea ice. Any associated northward movement of the southern
westerly wind belt would further reduce upwelling (Toggweiler
et al., 2006; Toggweiler, 2008, 2009).
Collectively, marine proxies indicate that the Southern Ocean
entered a fully stratified state as early as 70,000 years ago
(Robinson et al., 2007; Anderson et al., 2009), consistent with
increasing winter duration at that time. Opal records from the
Antarctic Polar Front (Anderson et al., 2009), as well as records of %
C. davisiana from the subantarctic zone (Hays et al., 1976), indicate
that the Southern Ocean then remained in a generally stratified
state, despite millennial-scale disturbances of stratification
between w60,000 and w36,000 yrs ago corresponding to Heinrich
Stadials in the North Atlantic (Anderson et al., 2009). These
disturbances were evidently unable to completely destratify the
ocean and initiate deep convection (Sigman et al., 2007). The
Southern Ocean entered a state of persistent full stratification
w36,000 yrs ago (Hays et al., 1976; Robinson et al., 2007; Anderson
et al., 2009; Sigman et al., 2010; Burke and Robinson, 2012), which
along with expanded winter sea ice around Antarctica (Gersonde
et al., 2003, 2005; Allen et al., 2011), continued through the LGM.
We suggest that stratification in the Southern Ocean, modulated
by the effects of Earth’s orbit on high-latitude winter duration
(Huybers and Denton, 2008), resulted in cold temperatures that
affected adjacent regions. Thus, Southern Ocean stratification
provides an explanation for the formation of Ohau glacier LLGM
moraines commencing as early as w32,500 yrs ago, and continuing
through to w18,200 yrs ago. Individual ice advance maxima of the
Ohau glacier during this interval may reflect millennial-scale fluctuations within generally cold climatic conditions.
A difficulty with the Southern Ocean coupled sea-ice/
stratification model is that, taken alone, the effects of winter
duration do not explain why, at w32,500 yrs ago, when southern
high-latitude winter duration was relatively short, the Southern
Ocean was stratified and the Ohau glacier was just as extensive as it
was w22,500 yrs ago, when winter duration was relatively long.
Although we consider southern winter duration to be an underlying factor governing stratification and SSTs in the Southern
Ocean, and hence southern mid-latitude glaciers, we note also that
Southern Alps glaciers would also have been strongly influenced by
Fig. 17. Plot comparing records mentioned in text. a.) 45! N (blue) and 45! S (red)
insolation intensity (Berger and Loutre, 1991). b.) Ohau glacier-inferred snowline and
temperature reconstruction. This curve incorporates the Lateglacial reconstruction
from Irishman Stream (Kaplan et al., 2010; Doughty et al., 2012). Solid black lines
depict where age control is known. Dotted black lines are based on inference. c.)
Patagonian glacier-inferred temperature changes (Denton et al., 1999a; Denton et al.,
1999b). Curve has been updated to reflect the Lateglacial record of Strelin et al.
(2011). d.) Modern analogue technique (MAT) SSTs and e.) G. ruber from core MD032611 south of Australia (De Deckker et al., 2012). The latter is a proxy for tropical
water incursions south of Australia, which indicates when the STF is to the south. f.)
South Atlantic SSTs (Barker et al., 2009). g.) EPICA Dronning Maud Land d18O (Antarctic
temperature proxy) (EPICA Community Members, 2006). Thin gray curve connects raw
data. Thick blue line is 10-pt moving average. h.) Atmospheric CO2 records of Monnin
et al. (2001), Indermühle et al. (1999, 2000), and Ahn and Brook (2008), synchronized
to GICC05 (Greenland methane) timescale (Barker et al., 2009). i.) Southern Ocean
stratification/upwelling record inferred from biogenic opal flux measured in cores
TN057-13PC and TN057-14PC (Anderson et al., 2009). Vertical yellow bands correspond with northern stadials (southern warming events) in Barker et al. (2009) and
Bond et al. (1999). HS1, HS2, and HS3 correspond to Heinrich Stadials 1, 2, and 3,
respectively. YD is Younger Dryas. ACR is Antarctic Cold Reversal. (For interpretation of
the references to colour in this figure legend, the reader is referred to the web version
of this article.)
Author's personal copy
138
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
the position of the STF upwind of South Island. Millennial-scale
migrations of the STF could have overwhelmed the more subtle
effects of winter duration at the northern edge of the Southern
Ocean between 36,000 and 18,000 yrs ago (De Deckker et al., 2012).
In addition, once stratified, the vertical structure of the Southern
Ocean, and its sea ice, could remain robust enough to withstand
disturbance by orbital forcing. Thus it seems that a more powerful
mechanism would have been required to destratify the Southern
Ocean and re-initiate deep ventilation (Sigman et al., 2007).
Before exploring mechanisms for disruption of Southern Ocean
stratification, we summarize the argument thus far. We speculate
that stratification in the Southern Ocean is driven fundamentally by
orbital forcing, as shown by the marine records of Hays et al. (1976),
in particular through the effect of high-latitude winter duration on
sea-ice formation (Huybers and Denton, 2008). Orbitally lengthened winters would prolong the period during which sea ice is
formed and exported to the open ocean by the circum-Antarctic
wind field, thereby promoting stratification in the open Southern
Ocean. Once formed, stratification appears difficult to disrupt. At
about 23,000 yrs ago, northern high-latitude summers were of low
insolation intensity and southern lobes of the Laurentide Ice Sheet
were at or near maximum extents (Lowell et al., 1999; Curry and
Grimley, 2006; Curry and Petras, 2011), while southern highlatitude winters were of long duration (Huybers and Denton,
2008; Huybers, 2009), and the Ohau glacier was at or near its
maximum extent. Thus, this mechanism, involving the effects of
southern winter duration on Southern Ocean sea ice and stratification, resolves Mercer’s (1984) conundrum, and offers an explanation for generally coeval maxima of the Ohau glacier and
northern ice sheets during the global LGM, despite opposing
summer insolation intensity patterns. This explanation circumvents any need for mysterious teleconnections between the
hemispheres in order to account for global synchrony of the LGM.
Future development of additional Southern Ocean stratification
records from marine proxies could help to test the links between
southern winter duration and stratification on orbital timescales.
We then come to the question of what brought about the end of
the southern LGM? Retreat of the Ohau glacier occurred in unison
with a regionally widespread warming trend, seen for example in
SSTs (Barrows et al., 2007; Calvo et al., 2007; Barker et al., 2009),
and indicated by disintegration of sea-ice cover around Antarctica
(Shemesh et al., 2002; Allen et al., 2011). This southern warming
coincided with Heinrich Stadial 1 (HS1) in the North Atlantic region
(Bond et al., 1999; Barker et al., 2009; Denton et al., 2010). Thus, at
face value, the southern warming beginning w18,000 yrs ago had
the hallmarks of an asynchronous millennial-scale climate oscillation between the North Atlantic and the Southern Ocean (Denton
et al., 2010), involving an oceanic (Crowley, 1992; Broecker, 1998)
and/or atmospheric (Toggweiler et al., 2006; Anderson et al., 2009)
bipolar seesaw mechanism between the hemispheres (Sigman
et al., 2007). However, on this occasion, a full glacial termination
resulted. Taken together, modeling of the Ohau glacier (this study)
and the Irishman Stream cirque glacier (Kaplan et al., 2010;
Doughty et al., 2012) shows an overall warming between the LLGM
and ACR of at least w3.6 ! C. Coeval phenomena included a net
poleward migration of the southern STF (Chiessi et al., 2008; Bard
and Rickaby, 2009; Sikes et al., 2009; De Deckker et al., 2012),
rising atmospheric CO2 (Monnin et al., 2001) and major upwelling
south of the Polar Front (Anderson et al., 2009), which indicates
breakdown of Antarctic stratification. Denton et al. (2010) pointed
out that poleward migration of the wind-driven southern STF is
difficult to explain solely by an oceanic mechanism, and proposed
that a southward shift of the southern westerlies at the onset of HS1
stimulated a poleward shift of the southern STF (Denton et al.,
2010; De Deckker et al., 2012). This southward shift of the
westerlies (Anderson et al., 2009), perhaps combined with the
effects of a bipolar seesaw operating through the interior of the
ocean (Broecker, 1998; Sigman et al., 2007), completely disrupted
the stratification of the glacial Southern Ocean, and thereby triggered the last mountain glacier termination in Southern Hemisphere middle latitudes.
In conclusion, we suggest that the ice-age signature as reflected
in the Ohau LLGM record, arose fundamentally from orbitally
driven changes in southern high-latitude winter duration that, in
turn, influenced stratification of the Southern Ocean. Previous
studies have established relationships between Southern Ocean
stratification and SSTs (Hays et al., 1976; Anderson et al., 2009). We
propose that a prolonged episode of Southern Ocean stratification
promoted generally cold southern mid-latitude atmospheric
temperatures, and hence full-glacial icefields, during the global
LGM. We then suggest that a bipolar seesaw oscillation, perhaps
operating through both the atmosphere and the ocean, caused
poleward migration of the southern STF and disruption of Southern
Ocean stratification beginning at 18,000 yrs ago, resulting in seasurface warming that produced the southern mid-latitude Last
Glacial termination.
8. Conclusions
1.) A precise and accurate 10Be surface-exposure chronology,
comprising 73 new dates on glacial landforms near Lake Ohau,
as well as 37 previously reported dates from the Irishman
Stream tributary of the Ohau catchment, indicates that the
LLGM in the Southern Alps was achieved as early as
32,520 " 970 yrs ago, and that subsequent glacier advances
produced slightly inboard moraine belts at 27,400 " 1300 yrs
ago, 22,510 " 660 yrs ago and 18,220 " 500 yrs ago. Recession
from the terminal moraine belt was underway by
17,690 " 350 yrs ago, and by 17,380 " 510 yrs ago, the Ohau
glacier tongue had retreated by as much as w24 km up valley.
This retreat represents a w40% decrease of the LLGM glacier
length and a net retreat rate of 77 m yr#1. A glacier that fed the
former Ohau glacier during the LLGM had receded well into the
mountains by 13,000 " 500 yrs ago (Kaplan et al., 2010).
2.) LLGM snowline was 920 " 50 m below the present day value
averaged across the Lake Ohau ice catchment. This snowline
depression corresponds to a temperature depression of about
6.25 " 0.5 ! C, which is the median value determined from
a variety of simulations in which flow parameters and precipitation were allowed to vary. Snowline and temperature rose by
at least w520 m and w3.6 ! C, respectively, between the LLGM
and ACR.
3.) Glacier variations at Lake Ohau complement marine and
terrestrial records of climate change in the southern middle
latitudes, as well as temperature and atmospheric CO2
concentrations measured in Antarctic ice cores. We propose
that enhanced sea ice and open-ocean stratification caused
a decrease in Southern Ocean SSTs and an equatorward shift in
the subtropical front that initiated the LLGM in the Southern
Alps. We think that this Southern Ocean signal could have been
paced in part by winter duration, which is sensitive to orbital
precession, thereby producing the LLGM at Lake Ohau coevally
with the maximum of the Laurentide Ice Sheet.
4.) We argue that a disruption of Southern Ocean stratification
beginning w18,000 yrs ago triggered the onset of the Last
Glacial termination in the Southern Alps. We attribute this
breakdown of stratification to a poleward shift of the southern
westerlies, associated with HS1 cooling in the North Atlantic
region (Denton et al., 2010). A southward shift of the southern
STF and invigorated Southern Ocean upwelling led to an abrupt
Author's personal copy
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
onset of sea-ice retreat, SST warming, and CO2 degassing,
thereby warming Southern Hemisphere atmospheric temperatures and inducing mountain glacier retreat at Lake Ohau in
the Southern Alps of New Zealand.
Acknowledgments
This work was supported by funding from the Gary C. Comer
Science and Education Foundation (CSEF), the Quesada Family
Foundation, the National Oceanographic and Atmospheric Administration (NOAA), and by National Science Foundation grants EAR1102782, EAR-0345835 and EAR-0745781. D.M. Sigman, S. Barker,
W.S. Broecker, R.F. Anderson, T.J. Hughes, P. Huybers, B.L. Hall,
G.R.M. Bromley, M.J. Vandergoes, P. Upton, and R.B. Alley provided
helpful discussions. In particular, we are grateful to J.D. Hays for
engaging discussions about the role of polar ocean stratification in
ice-age cycles. N.R. Golledge, A.N. Mackintosh, B. Anderson, and
T.J.H. Chinn contributed helpful insights into glaciological modeling
in the Southern Alps. S. Barker kindly provided data from core
TNO57-21. B. Lemieux-Dudon provided the updated EPICA Dome C
chronology. K. Ladig and S. Travis helped with fieldwork. J. Frisch
assisted with laboratory work. We are grateful to the Kees family
(Ohau Station), F. Hawkins, M. Lindsay, J. Papich, G. Burrows (Ribbonwood Station), S. and H. Williamson (Glenbrook Station), J. and
K. Wigley (Glen Lyon Station), and the Department of Conservation
! nanga o Nga
!i Tahu, for permitting us
e Te Papa Atawhai, Te Ru
access to their land. T. and K. Ritchie of Lake Ruataniwha Holiday
Park provided excellent accommodation during long field seasons.
V. Jomelli and an anonymous reviewer provided helpful suggestions that improved the paper. A.E. Putnam was supported by CSEF,
NOAA, a University of Maine teaching assistantship, and a Lamonte
Doherty Earth Observatory (LDEO) postdoctoral fellowship while
conducting this research. GNS Science’s Direct Crown Funded
Programme ‘Global Change through Time’ supported D.J.A.B. This is
LDEO contribution #7629.
Appendix A
A complete description of UMISM can be found in Fastook and
Prentice (1994) and Fastook et al. (2008). Here we provide an
overview of the model ice-flow solver in order to give context for
the tuning reported in the methods section. The tuning pertains to
mainly the sliding and ice hardness terms below in Equations (2),
(3) and (5).
The core of UMISM is a differential equation for ice extent and
thickness as a function of time derived from an integrated
momentum conservation equation based on the flow law of ice
(Glen, 1955), coupled with a continuity equation for mass
conservation:
dh
¼ MB # V$ðUHÞ;
dt
(1)
where dh/dt is the time-dependent ice-surface elevation, MB is the
net-annual surface mass balance (accumulation minus ablation), U
is the column-average ice velocity, and H is the ice thickness.
U in Equation (1) is obtained by expanding:
3_
¼
h s in
A
;
(2)
which is a tensor equation expressing strain rates, 3_ , and stresses, s,
through a non-linear power law. In this equation, n is the empirical
flow law constant [we use a value of 3 for temperate glaciers
139
(Paterson, 1994)], and A is a temperature dependent function that
represents ice hardness in an Arrhenius relationship:
A ¼ EA0 e#Q =RT ;
(3)
where A0 is a constant, Q is the activation energy for ice creep, R is
the gas constant, T is the ice temperature, and E is a tuning (or flow
enhancement) parameter meant to account for ice impurities.
By shallow-ice approximation, all stresses and strain rates are
ignored in the force balance except basal drag. Thus, the stress in
Equation (2) acting on the bed, or the “driving” stress, is expressed:
(4)
sxz ¼ rgHVh;
where sxz is a stress in the x-direction acting on a surface with
normal in the z direction, r is density of ice, g is gravitational
acceleration, H is ice thickness, and Vh is ice-surface slope.
Finally, U, is calculated from the sum of velocity due to internal
deformation, UD, and velocity due to basal sliding, US, as expressed:
U ¼ UD þ US ¼
"
"
#
#
rgVh m m
2 rgVh n nþ1
H
þ
H ;
nþ2 A
B
(5)
where the last term follows a general sliding law relationship
developed by Weertman (1964) for beds at the melting point, and
modified to incorporate the effect of basal water (Johnson and
Fastook, 2002). Here n is the empirical flow law constant in Equation (2), A is the ice hardness parameter described in Equation (3), B
is the sliding law constant, and m is the sliding law exponent.
References
Ahn, J., Brook, E.J., 2008. Atmospheric CO2 and climate on millennial time scales
during the last glacial period. Science 322 (5898), 83e85.
Allen, C.S., Pike, J., Pudsey, C.J., 2011. Last GlacialeInterglacial sea-ice cover in the
SW Atlantic and its potential role in global deglaciation. Quaternary Science
Reviews 19e20, 2446e2458.
Amos, C.B., Burbank, D.W., Nobes, D.C., Read, S.A.L., 2007. Geomorphic constraints
on listric thrust faulting: Implications for active deformation in the Mackenzie
Basin, South Island, New Zealand. Journal of Geophysical Research 112. http://
dx.doi.org/10.1029/2006JB004291.
Amos, C.B., Burbank, D.W., Read, S.A.L., 2010. Along-strike growth of the Ostler fault,
New Zealand: consequences for drainage deflection above active thrusts.
Tectonics 29, TC4021.
Anderson, B., Mackintosh, A., 2006. Temperature change is the major driver of lateglacial and Holocene glacier fluctuations in New Zealand. Geology 34 (2),121e124.
Anderson, B., Lawson, W., Owens, I., Goodsell, B., 2006. Past and future mass balance
of ‘Ka Roimata o Hine Hukatere’ Franz Josef Glacier, New Zealand. Journal of
Glaciology 52 (179), 597e607.
Anderson, R.F., et al., 2009. Wind-driven upwelling in the Southern Ocean and the
deglacial rise in atmospheric CO2. Science 323, 1443e1448.
Anderson, B., et al., 2010. Climate sensitivity of a high-precipitation glacier in New
Zealand. Journal of Glaciology 56 (195), 114e128.
Applegate, P.J., et al., 2012. Improved moraine age interpretations through explicit
matching of geomorphic process models to cosmogenic nuclide measurements
from single landforms. Quaternary Research 77 (2), 293e304.
Augustinus, P.C., D’Costa, D., Deng, Y., Hagg, J., Shane, P., 2011. A multi-proxy record
of changing environments from ca. 30 000 to 9000 cal. a BP: Onepoto maar
palaeolake, Auckland, New Zealand. Journal of Quaternary Science 26 (4),
389e401.
Balco, G., Schaefer, J.M., 2006. Cosmogenic-nuclide and varve chronologies for the
deglaciation of southern New England. Quaternary Geochronology 1 (1), 15e28.
Balco, G., Stone, J.O., Lifton, N.A., Dunai, T.J., 2008. A complete and easily accessible
means of calculating surface exposure ages or erosion rates from 10Be and 26Al
measurements. Quaternary Geochronology 4, 93e107.
Bard, E., Rickaby, R.E.M., 2009. Migration of the subtropical front as a modulator of
glacial climate. Nature 460, 380e384.
Barker, S., et al., 2009. Interhemispheric Atlantic seesaw response during the last
deglaciation. Nature 457 (7233), 1097e1102.
Barrell, D.J.A., Andersen, B.G., Denton, G.H., 2011. Glacial geomorphology of the
central South Island, New Zealand. GNS Science Monograph 27. Lower Hutt,
New Zealand. Map and 71 pp.
Barrell, D.J.A., 2011. Quaternary glaciers of New Zealand. In: Ehlers, J., Gibbard, P.L.,
Hughes, P.D. (Eds.), Quaternary Glaciations e Extent and Chronology, Part IV e
a Closer Look. Developments in Quaternary Science. Elsevier, Amsterdam,
pp. 1047e1064.
Author's personal copy
140
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
Barrows, T.T., Juggins, S., De Deckker, P., Calvo, E., Pelejero, C., 2007. Long-term sea
surface temperature and climate change in the Australian-New Zealand region.
Paleoceanography 22, PA22215.
Berger, A., Loutre, M.F., 1991. Insolation values for the climate of the last 10000000
years. Quaternary Science Reviews 10 (4), 297e317.
Bevington, P., Robinson, D., 1992. Data Reduction and Error Analysis for the Physical
Sciences. WCB McGraw Hill.
Birkel, S.D., et al., 2012. Climate inferences from a glaciological reconstruction of the
Late Pleistocene Wind River Ice Cap, Wind River Range, Wyoming. Arctic,
Antarctic, and Alpine Research 44 (3), 265e276.
Birkeland, P.W., 1982. Subdivision of Holocene glacial deposits, Ben Ohau Range,
New Zealand, using relative-dating methods. Geological Society of America
Bulletin 93, 443e449.
Blunier, T., Brook, E.J., 2001. Timing of millennial-scale climate change in Antarctica
and Greenland during the last glacial period. Science 291 (5501), 109e112.
Bond, G.C., et al., 1999. The north Atlantic’s 1e2 kyr climate rhythm: relation to
Heinrich events, Dansgaard/oeschger Cycles and the Little Ice Age. Geophysical
Monograph 112, 35e58.
Bouttes, N., Paillard, D., Roche, D.M., 2010. Impact of brine-induced stratification on
the glacial cycle. Climate of the Past Discussions 6, 575e589.
Bouttes, N., et al., 2011. Impact of oceanic processes on the carbon cycle during the
last termination. Climate of the Past Discussions 7, 1887e1934.
Braithwaite, R.J., Olesen, O.B., 1989. Calculation of glacier ablation from air
temperature, West Greenland. In: Oerlemans, J. (Ed.), Glacier Fluctuations. D.
Reidel Publishing Company, Dordrecht, p. 160.
Braithwaite, R.J., 2011. Why do we expect glacier melting to increase under global
warming? In: Blanco, J., Kheradman, H. (Eds.), Climate Change e Geophysical
Foundations and Ecological Effects. InTech, p. 520.
Broecker, W.S., van Donk, J., 1970. Insolation changes, ice volume, and the 18O record
in deep-sea cores. Reviews of Geophysics 8 (1), 169e198.
Broecker, W.S., 1966. Absolute dating and the astronomical theory of glaciation.
Science 151, 299e304.
Broecker, W.S., 1998. Paleocean circulation during the last deglaciation: a bipolar
seesaw? Paleoceanography 13 (2), 119e121.
Brown, E., et al., 2001. Ocean Circulation, second ed. The Open University.
Burke, A., Robinson, L.F., 2012. The Southern Ocean’s role in carbon exchange during
the last deglaciation. Science 335, 557e561.
Burrows, C.J., 2005. Julius Haast in the Southern Alps. Canterbury University Press,
Christchurch.
Calvo, E., Pelejero, C., De Deckker, P., Logan, G.A., 2007. Antarctic deglacial pattern in
a 30 kyr record of sea surface temperature offshore South Australia. Geophysical
Research Letters 34 (13), L13707.
Carter, L., et al., 1998. Ocean Circulation New Zealand. In: NIWA Chart Miscellaneous
Series, vol. 76. NIWA, Wellington.
Chiessi, C.M., et al., 2008. South Atlantic interocean exchange as the trigger for the
Bolling warm event. Geology 36 (12), 919e922.
Chinn, T., Winkler, S., Salinger, M.J., Haakensen, N., 2005. Recent glacier advances in
Norway and New Zealand: a comparison of their glaciological and meterological causes. Geografiska Annaler 87A, 141e157.
Chinn, T.J., 1996. New Zealand glacier responses to climate change of the past
century. New Zealand Journal of Geology and Geophysics 39, 415e428.
Clark, P.U., Mix, A.C., 2002. Ice Sheets and sea level of the Last Glacial Maximum.
Quaternary Science Reviews 21, 1e7.
Clark, P.U., et al., 2009. The Last Glacial Maximum. Science 325, 710e714.
Cox, S.C., Barrell, D.J.A., 2007. Geology of the Aoraki Area. 1:250,000 Geological Map
15. GNS Science. GNS Science, Lower Hutt, New Zealand, 1 Sheet and 71 pp.
Crowley, T.J., 1992. North Atlantic deep water cools the Southern Hemisphere.
Paleoceanography 7 (4), 489e497.
Curry, B.B., Grimley, D.A., 2006. Provenance, age, and environment of midWisconsinan slackwater lake sediment in the St. Louis Metro East area, USA.
Quaternary Research 65, 108e122.
Curry, B.B., Petras, J., 2011. Chronological framework for the deglaciation of the Lake
Michigan lobe of the Laurentide Ice Sheet from ice-walled deposits. Journal of
Quaternary Science 26 (4), 402e410.
Davis, K., Burbank, D.W., Fisher, D., Wallace, S., Nobes, D., 2005. Thrust-fault growth
and segment linkage in the active Ostler fault zone, New Zealand. Journal of
Structural Geology 27 (8), 1528e1546.
De Deckker, P., Moros, M., Perner, K., Jansen, E., 2012. Influence of the tropics and
southern westerlies on glacial interhemispheric asymmetry. Nature Geoscience
5, 266e269.
Dee, D.P., et al., 2011. The ERA-Interim reanalysis: configuration and performance of
the data assimilation system. Quarterly Journal of the Royal Meteorological
Society 137, 553e597.
Denton, G.H., et al., 1999a. Interhemispheric linkage of paleoclimate during the last
glaciation. Geografiska Annaler 81A (2), 107e153.
Denton, G.H., et al., 1999b. Geomorphology, stratigraphy, and radiocarbon chronology of Llanquihue drift in the area of the southern Lake District, Seno
Reloncaví, and Isla Grande de Chiloé, Chile. Geografiska Annaler 81A, 167e229.
Denton, G.H., et al., 2010. The last glacial termination. Science 328, 1652e1656.
Desilets, D., Zreda, M., Prabu, T., 2006. Extended scaling factors for in situ cosmogenic nuclides: new measurements at low latitude. Earth and Planetary Science
Letters 246, 256e276.
Doughty, A.M., et al., 2012. Late-glacial temperatures in the Southern Alps of New
Zealand based on glacier modelling at Irishman basin. Quaternary Science
Reviews. http://dx.doi.org/10.1016/j.quascirev.2012.09.013.
Drost, F., Renwick, J., Bhaskaran, B., Oliver, H., McGregor, J., 2007. A simulation of
New Zealand’s climate during the Last Glacial Maximum. Quaternary Science
Reviews 26, 2505e2525.
Dunai, T., 2001. Influence of secular variation of the magnetic field on production
rates of in situ produced cosmogenic nuclides. Earth and Planetary Science
Letters 193, 197e212.
Dunai, T.J., 2010. Cosmogenic Nuclides: Principles, Concepts, and Applications in the
Earth Surface Sciences. Cambridge University Press, New York.
EPICA Community Members, 2006. One-to-one coupling of glacial climate variability in Greenland and Antarctica. Nature 444 (7116), 195e198.
Fastook, J.L., Prentice, M., 1994. A finite-element model of Antarctica: sensitivity test
for meteorological mass-balance relationship. Journal of Glaciology 40 (134),
167e175.
Fastook, J.L., Head, J.W., Marchant, D.R., Forget, F., 2008. Tropical mountain glaciers
on Mars: altitude-dependence of ice accumulation, accumulation conditions,
formation times, glacier dynamics, and implications for planetary spin-axis/
orbital history. Icarus 198, 305e317.
Gair, H.S., 1967. Sheet 20-Mount Cook, Geological Map of New Zealand. Department
of Scientific and Industrial Research, Wellington.
Gersonde, R., et al., 2003. Last glacial sea surface temperatures and sea-ice extent in
the Southern Ocean (Atlantic-Indian sector): a multiproxy approach. Paleoceanography 18 (3). http://dx.doi.org/10.1029/2002PA000809.
Gersonde, R., Crosta, X., Abelmann, A., Armand, L., 2005. Sea-surface temperature
and sea ice distribution of the Southern Ocean at the EPILOG Last Glacial
Maximum e a circum-Antarctic view based on siliceous microfossil records.
Quaternary Science Reviews 24 (7e9), 869e896.
Ghisetti, F.C., Gorman, A.R., Sibson, R.H., 2007. Surface breakthrough of a basement
fault by repeated seismic slip episodes: the Ostler Fault, South Island, New
Zealand. Tectonics 26. http://dx.doi.org/10.1029/2007TC002146.
Glen, J.W., 1955. The creep of polycrystalline ice. Proceedings of the Royal Society of
London 228, 519e538.
Golledge, N.R., et al., 2012. Last Glacial Maximum climate in New Zealand inferred
from a modelled Southern Alps icefield. Quaternary Science Reviews 46, 30e45.
Gordon, A.L., 1981. Seasonality of Southern Ocean sea ice. Journal of Geophysical
Research 86, 4193e4197.
Gordon, A.L., 1991. Two stable modes of Southern Ocean winter stratification. In:
Chu, P.C., Gascard, J.C. (Eds.), Deep Convection and Deep Water Formation in the
Oceans. Elsevier Science Publishers B. V..
Hays, J.D., Imbrie, J., Shackleton, N.J., 1976. Variations in the earth’s orbit: pacemaker
of the ice ages. Science 194 (4270), 1121e1132.
Henderson, R.D., Thompson, S.M., 1999. Extreme rainfalls in the Southern Alps of
New Zealand. Journal of Hydrology (NZ) 38, 309e330.
Heusser, C.J., 1990. Chilotan piedmont glacier in the southern Andes during the last
glacial maximum. Revista Geologíca de Chile 17, 3e18.
Hijmans, R., Cameron, S.E., Parra, J.L., Jones, P.G., Jarvis, A., 2005. Very high resolution interpolated climate surfaces for global land areas. International Journal
of Climatology 25, 1965e1978.
Huybers, P., Denton, G., 2008. Antarctic temperature at orbital timescales controlled
by local summer duration. Nature Geoscience 1 (11), 787e792.
Huybers, P., 2009. Antarctica’s orbital beat. Science 325 (5944), 1085e1086.
Indermuhle, A., et al., 1999. Holocene carbon-cycle dynamics based on CO2 trapped
in ice at Taylor Dome, Antarctica. Nature 398 (6723), 121e126.
Indermühle, A., Monnin, E., Stauffer, B., Stocker, T.F., Wahlen, M., 2000. Atmospheric
CO2 concentration from 60 to 20 kyr BP from the Taylor Dome ice core,
Antarctica. Geophysical Research Letters 27 (5), 735e738.
Johnson, J., Fastook, J.L., 2002. Northern Hemisphere glaciations and its sensitivity
to basal melt water. Quaternary International 95e96, 65e74.
Kaiser, J., Lamy, F., Hebbeln, D., 2005. A 70-kyr sea surface temperature record off
southern Chile (Ocean Drilling Program Site 1233). Paleoceanography 20 (4),
PA4009. http://dx.doi.org/10.1029/2005PA001146.
Kaplan, M.R., et al., 2010. Glacier retreat in New Zealand during the Younger Dryas
stadial. Nature 467, 194e197.
Kaplan, M.R., et al., 2011. In-situ cosmogenic 10Be production rate at Lago Argentino,
Patagonia: implications for late-glacial climate chronology. Earth and Planetary
Science Letters 309, 21e32.
Kawamura, K., et al., 2007. Northern Hemisphere forcing of climatic cycles in
Antarctica over the past 360,000 years. Nature 448 (7156), 912e916.
Keeling, R.F., Visbeck, M., 2011. On the linkage between Antarctic surface water
stratification and global deep-water temperature. Journal of Climate 24,
3545e3557.
Kelly, M.A., 2003. The late Würmian Age in the western Swiss Alps e last glacial
maximum (LGM) ice-surface reconstruction and 10Be dating of late-glacial
features. Ph.D. thesis. University of Bern, 105 pp.
Laabs, B.J.C., Plummer, M.A., Mickelson, D.M., 2006. Climate during the last glacial
maximum in the Wasatch and southern Uinta Mountains inferred from glacier
modeling. Geomorphology 75, 300e317.
Labeyrie, L., et al., 1996. Hydrographic changes of the Southern Ocean (southeast
Indian sector) over the last 230 kyr. Paleoceanography 11 (1), 57e76.
Lamont, G.N., Chinn, T.J.H., Fitzharris, B.B., 1999. Slopes of glacier ELAs in the
Southern Alps of New Zealand in relation to atmospheric circulation patterns.
Global and Planetary Change 22, 209e219.
Lamy, F., et al., 2004. Antarctic timing of surface water changes off Chile and
Patagonian ice sheet response. Science 304 (5679), 1959e1962.
Lamy, F., et al., 2007. Modulation of the bipolar seesaw in the southeast pacific
during Termination 1. Earth and Planetary Science Letters 259 (3e4), 400e413.
Author's personal copy
A.E. Putnam et al. / Quaternary Science Reviews 62 (2013) 114e141
Le Meur, E., Vincent, C., 2003. A two-dimensional shallow ice flow model of glacier
de Saint Sorlin, France. Journal of Glaciology 49, 527e538.
Le Meur, E., Gagliardini, O., Zwinger, T., Ruokolainen, J., 2004. Glacier flow modelling: a comparison of the shallow ice approximation and the full-Stokes solution. Comptes Rendus Physique 5 (7), 709e722.
Le Meur, E., Gerbaux, M., Schäfer, M., Vincent, C., 2007. Disappearance of an alpine
glacier over the 21st century simulated from modeling its future surface mass
balance. Earth and Planetary Science Letters 261, 367e374.
Lemieux-Dudon, B., et al., 2010. Consistent dating for Antarctic and Greenland ice
cores. Quaternary Science Reviews 29 (1e2), 8e20.
Leysinger Vieli, G., Gudmundsson, G.H., 2004. On estimating length fluctuations of
glaciers caused by changes in climatic forcing. Journal of Geophysical Research
109 (F1). http://dx.doi.org/10.1029/2003JF000027.
Lifton, N., et al., 2005. Addressing solar modulation and long-term uncertainties in
scaling secondary cosmic rays for in situ cosmogenic nuclide applications. Earth
and Planetary Science Letters 239, 140e161.
Lifton, N., Smart, B., Shea, M., 2008. Scaling time-integrated in situ cosmogenic
nuclide production rates using a continuous geomagnetic model. Earth and
Planetary Science Letters 268, 190e201.
Lowell, T.V., Hayward, R.K., Denton, G.H., 1999. Role of climate oscillations in
determining ice-margin position: hypothesis, examples, and implications.
Geological Society of America Special Papers 337, 193e203.
MARGO Project Members, 2009. Constraints on the magnitude and patterns of
ocean cooling at the Last Glacial Maximum. Nature Geoscience 2, 127e132.
McGlone, M.S., Moar, N.T., 1998. Dryland Holocene vegetation history, Central Otago
and the Mackenzie Basin, South Island, New Zealand. New Zealand Journal of
Botany 36, 91e111.
Mercer, J.H., 1972. Chilean glacial chronology 20,000 to 11,000 carbon-14 years ago:
some global comparisons. Science 176, 1118e1120.
Mercer, J.H., 1984. Simultaneous climatic change in both hemispheres and similar
bipolar warming: evidence and implications. In: Hansen, J.E., Takahashi, T.
(Eds.), Climate Processes and Climate Sensitivity. Geophysical Monograph
American Geophysical Union, Washington. D. C., pp. 307e313.
Milankovitch, M., 1941. Kanon der Erdbestrahlung und seine Andwendung auf das
Eiszeiten-problem. Royal Serbian Academy, Belgrade.
Mix, A.C., Bard, E., Schneider, R., 2001. Environmental processes of the ice age: land,
oceans, glaciers (EPILOG). Quaternary Science Reviews 20, 627e657.
Monnin, E., et al., 2001. Atmospheric CO2 concentrations over the last glacial
termination. Science 291, 112e114.
Newnham, R.M., Loewe, D.J., Giles, T., Alloway, B.V., 2007a. Vegetation and climate
of Auckland, New Zealand, since ca. 32,000 cal. yr ago: support for an extended
LGM. Journal of Quaternary Science 22 (5), 517e534.
Newnham, R.M., Vandergoes, M.J., Hendy, C.H., Lowe, D.J., Preusser, F., 2007b.
A terrestrial palynological record for the last two glacial cycles from southwestern New Zealand. Quaternary Science Reviews 26, 517e535.
Nishiizumi, K., et al., 2007. Absolute calibration of 10Be AMS standards. Nuclear
Instruments and Methods in Physics Research B 258, 403e413.
Norton, D., 1985. A multivariate technique for estimating New Zealand temperature
normals. Weather and Climate 5, 64e74.
Oerlemans, J., 1997. Climate sensitivity of Franz Josef Glacier, New Zealand, as
revealed by numerical modelling. Arctic and Alpine Research 29 (2), 233e239.
Paterson, W.S.B., 1994. The Physics of Glaciers. Pergammon, Oxford.
Pelejero, C., Calvo, E., Barrows, T.T., Logan, G.A., De Deckker, P., 2006. South Tasman
Sea alkenone palaeothermometry over the last four glacial/interglacial cycles.
Marine Geology 230, 73e86.
Petit, J.R., et al., 1999. Climate and atmospheric history of the past 420,000 years
from the Vostok ice core, Antarctica. Nature 399 (6735), 429e436.
Porter, S.C., 1975. Equilibrium-line altitudes of Late Quaternary glaciers in the
Southern Alps, New Zealand. Quaternary Research 5, 27e47.
Purdie, H., et al., 2011. Interannual variability in net accumulation on Tasman Glacier
and its relationship with climate. Global and Planetary Change. http://
dx.doi.org/10.1016/j.gloplacha.2011.04.004.
Putkonen, J., Swanson, T., 2003. Accuracy of cosmogenic ages for moraines.
Quaternary Research 59, 255e261.
Putnam, A.E., et al., 2010a. Glacier advance in southern middle latitudes during the
Antarctic Cold Reversal. Nature Geoscience 3, 700e704.
141
Putnam, A.E., et al., 2010b. In situ cosmogenic 10Be production-rate calibration from
the Southern Alps, New Zealand. Quaternary Geochronology 5, 392e409.
Refsnider, K.A., et al., 2008. Last glacial maximum climate inferences from cosmogenic dating and glacier modeling of the western Uinta ice field, Uinta Mountains, Utah. Quaternary Research 69, 130e144.
Reimer, P.J., et al., 2009. INTCAL09 and MARINE09 radiocarbon age calibration
curves, 0e50,000 years cal BP. Radiocarbon 51 (4), 1111e1150.
Robinson, R.S., Mix, A., Martinez, P., 2007. Southern Ocean control on the extent of
denitrification in the southeast Pacific over the last 70 ka. Quaternary Science
Reviews 26 (1e2), 201e212.
Roe, G., 2006. In defense of Milankovitch. Geophysical Research Letters 33 (L24703).
http://dx.doi.org/10.1029/2006GL027817.
Sandiford, A., Newnham, R.M., Alloway, B.V., Ogden, J., 2003. A 28 000e7600 yr BP
pollen record of vegetation and climate change from the Pukaki Crater,
northern New Zealand. Palaeogeography Palaeoclimatology Palaeoecology 201,
235e247.
Schaefer, J.M., et al., 2006. Near-synchronous interhemispheric termination of the
last glacial maximum in mid-latitudes. Science 312 (5779), 1510e1513.
Schaefer, J.M., et al., 2009. High-frequency Holocene glacier fluctuations in New
Zealand differ from the northern signature. Science 324, 622e625.
Schäfer, M., Gagliardini, O., Pattyn, F., Le Meur, E., 2008. Applicability of the shallow
ice approximation inferred from model inter-comparison using various glacier
geometries. Cryosphere Discussions 2 (4), 557e599.
Shemesh, A., et al., 2002. Sequence of events during the last deglaciation in
Southern Ocean sediments and Antarctic ice cores. Paleoceanography 17 (4).
http://dx.doi.org/10.1029/2000PA000599.
Sigman, D.M., Haug, G.H. (Eds.), 2003. The Biological Pump in the Past. Treatise on
Geochemistry, vol. 6. Elsevier Pergamon.
Sigman, D.M., Jaccard, S.L., Haug, G.H., 2004. Polar ocean stratification in a cold
climate. Nature 428 (6978), 59e63.
Sigman, D.M., de Boer, A.M., Haug, G.H., 2007. Antarctic stratification, atmospheric
water vapor, and Heinrich Events: a hypothesis for Late Pleistocene deglaciations. In: Schmittner, A., Chiang, J.C.H., Hemming, S.R. (Eds.), Ocean Circulation:
Mechanisms and Impacts. Geophysical Monograph Series, vol. 173. American
Geophysical Union, Washington, pp. 335e349.
Sigman, D.M., Hain, M.P., Haug, G.H., 2010. The polar ocean and glacial cycles in
atmospheric CO2 concentration. Nature 466, 47e55.
Sikes, E.L., et al., 2009. Southern Ocean seasonal temperature and Subtropical Front
movement on the South Tasman Rise in the late Quaternary. Paleoceanography
24, PA2201.
Staiger, J., et al., 2007. Atmospheric scaling of cosmogenic nuclide production:
climate effect. Journal of Geophysical Research 112. http://dx.doi.org/10.1029/
2005JB003811.
Stone, J.O., 2000. Air pressure and cosmogenic isotope production. Journal of
Geophysical Research 105, 23753e23759.
Strelin, J.A., Denton, G.H., Vandergoes, M.J., Ninnemann, U.S., Putnam, A.E., 2011. Radiocarbon chronology of the late-glacial Puerto Bandera moraines, Southern Patagonian Icefield, Argentina. Quaternary Science Reviews 30 (19e20), 2551e2569.
Sugden, D.E., John, B.S., 1976. Glaciers and Landscape: a Geomorphological
Approach. Edward Arnold, London.
Suggate, R.P., 1990. Late Pliocene and Quaternary glaciations of New Zealand.
Quaternary Science Reviews 9, 175e197.
Toggweiler, J.R., Russell, J.L., Carson, S.R., 2006. Midlatitude westerlies, atmospheric
CO2, and climate change during the ice ages. Paleoceanography 21 (2), PA2005.
Toggweiler, J.R., 2008. Origin of the 100,000-year timescale in Antarctic temperatures and atmospheric CO2. Paleoceanography 23 (2), PA2211. http://dx.doi.org/
10.1029/2006PA001405.
Toggweiler, J.R., 2009. Shifting westerlies. Science 323 (5920), 1434e1435.
Vandergoes, M.J., et al., 2005. Regional insolation forcing of late Quaternary climate
change in the Southern Hemisphere. Nature 436, 242e245.
Ward, G.K., Wilson, S.R., 1978. Procedures for comparing and combining radiocarbon age determinations: a critique. Archaeometry 20, 19e31.
Weertman, J., 1964. The theory of glacier sliding. Journal of Glaciology 5 (39), 287e303.
Whittaker, T.E., Hendy, C.H., Hellstrom, J.C., 2011. Abrupt millennial-scale changes in
intensity of Southern Hemisphere westerly winds during marine isotope stages
2e4. Geology 39 (5), 455e458.