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Study of Glacier flow Glacier Flow and Ice Profiles •  Based on three lines of observaMon: 1.  Field evidence 2.  Laboratory experiments 3.  TheoreMcal arguments What can we learn about glacier movement and subglacial processes? UMass Geo 563, Fall 2009, 1&6 October •  Defining the deformaMon, viscous flow, and sliding of glacial ice under the force of gravity. Svalbard, 2005 Glaciers transfer mass from
the Accumulation zone to
the ablation zone.
Downslope transport of ice mass Most rapid velocities
reached at the ELA
Accumula'on vs abla'on Why?
ELA
Glaciers maintain their shape by flowing • to carry away what has accumulated • necessary to adjust their shape to maintain flow Ice has rheological properMes (sMffness) Flow is determined by forces (stresses) applied Accum Area dominated
by extention, with
deformation driving ice
toward the bed of the
glacier;
Ablation area
dominated by
compressional flow.
Cartoons from www.physicalgeography.net 1
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Zone of extension – driving particles downward
ice
Zone of compression – driving particles upward
ice
Cartoons from www.physicalgeography.net Ed Waddington, notes
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Field observations of
motion:
Complex
crevasses
shown on an
active glacier.
Note lower
figure with
strain ellipses
Fractures
perpendicular
to strain.
Ed Waddington, notes
Can use inSAR
methods to infer ice
velocities and strain
rates
Flow varies with depth over time
Worthington Glacier
Harper et al., 1998 Science
inSAR =interferometric
Synthetic Aperature
Radar
Example from Ross
Ice Shelf – Joughin
and Tulaczyk, 2002,
Science.
Crevasses can be
tracked with repeated
imagery.
Scale is 0 to 800 meters / yr
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What determines Glacier motion?
Ice DeformaMon (creep) • Ice flow law • Temperature • Fabric, grain size • Chemistry, impuriMes ions, bubbles, rocks parMcles Requires: Viscous flow, ElasMc response Brible behavior MoMon over the Bed • Basal sliding ‐‐requires water • Bed deformaMon ‐‐requires deformable substrate Movement at the surface (Us) is the accumulaMve effect of these processes, acMng singularly or in combinaMon! Avalanching • Fracture / failure Ed Waddington, notes
Ed Waddington, notes
Stress ‐‐ Force per unit area ice in glacier subjected to weight of ice and the slope of of the ice surface Strain – change in shape of a material due to Stress. Ed
Waddingto
nnotes
Force = Mass x gravitational acceleration
- elastic response
recoverable deformation
- permanent deformation
brittle or ductile change
with flow or creep
Strain Rate = a‐b/a = x/yr Where a= original length b= remeasured length afer some interval of Mme 4
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Ice Deforms as a non-linear
viscous material
Creep behavior of ice Creep – movements
within or between
grains;
May involve gliding
by cleavage planes;
recrystalization; and
dislocations within
crystalline lattice of
the ice.
See
http://
web.earthsci.unimelb.edu.
au/wilson/ice1/index.html
For movies of ice
deformation
Primary = strain rate decreases over time; grain hardening
Secondary = steady state creep at minimum strain rate
Tertiary = increasing creep with strain
Final Steady state creep --
Ed Waddington, notes
Glen’s Flow Law •  Simple power relaMonship, where strain εs is described as : εs = B σ n Where B = temperature dependent ice hardness parameter 0 deg C B=0.165 ‐5 deg C B = 0.054 ‐10 deg C B= 0.017 ‐ 20 deg C b= 0.0016 σ = amt of Stress n= creep exponent (1=linear flow; ∞ = perfect plasMc but usually between 1.9 and 4.5 with a mean of 3) Cold ice more rigid and behaves slowly; temperate glacier ice deforms faster than polar glacier. Ed Waddington, notes
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To model ice sheet flow think of ice as a slab spreading by creep under its own wt. α
h
Τau = ρ g h (sin α)
Tau for most glaciers………
= 0.5 to 1.0 bars
= 50 to 100 kPa
With 1 a best estimate close to yield
strength of ice.
BUT!
For fast flow ice streams and outlet
glaciers can be as low as 4-20 kPa
Basal shear stress = ice density x gravitation acceleration x thickness x
sin of the surface slope; Ice always flows in direction of surface slope.
To compute ice thickness, =h=
Tau
ρg(sin α)
What does this
mean?
Glacier bed condiMons Normal Stress acMng in a verMcal plane and no water present: EffecMve Stress for situaMons with water at the base: δi = ρi g hi
δe = ρi g hi - pw
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