Sea-level, Ice volume, and the O- isotopic composition of the ocean

Sea-level, Ice volume, and the Oisotopic composition of the
ocean
Warm
Cold
Low elevation
Low latitude
High elevation
High latitude
-34‰
δ O
-26‰
-16‰
-10‰
-20‰
-12‰
-4‰
-10‰
0‰
0‰
Rayleigh Distillation: R = R0fα -1
slide from Clark and Aravena
T-dependence δ 18O of calcification
Ca16CO3 + H218O ↔ CaC16O218O + H216O
T = 16.9 – 4.2(δ18Oc – δ18Ow)
Epstein et al., 1953
δ18 Ow = f(precipitation, evaporation, runoff, global ice volume)
1
Variability of δ 18O at the surface
Variability
determined by
evaporation and
precipitation/runoff
Range of δ18O in
surface is as large
as the signal
imprinted by ice
volume changes
on the glacialinterglacial time
scale.
From “ The Glacial World According to Wally ”
Sea-level/Ice volume Proxy: δ 18O of
Benthic Foraminifera
Benthic forams live in the deep sea, where the δ18O composition is
relatively homogenous compared to surface water
∆δ18Oglacial-interglacial ˜ 1.7 ‰
from “ The Glacial World According to Wally ”
δ18 O of carbonates in
benthic forams reflect
the isotopic
composition of seawater and
temperature.
Assumed to be
mostly a seawater
δ18 O affect, since
temperature of deep
seawater isn’t
expected to vary
much with glacialinterglacial cycles.
2
Separation of ∆T and ∆δ18Osw from
marine isotope record
• Constant temperature assumption – in regions where past
temperature change is assumed to be minimal (e.g. Shackelton,
1967) --> because bottom water is currently only a few degrees
above its freezing point
• Mass balance calculations (e.g. Broecker, 1975)
• Porewater δ18O (Schrag and DePaolo, 1993)
• δ18O of air O2 from ice cores (Shackelton, 2000)
• Others, such as coral δ18O (discussed by Julian)
Mass balance calculation with
constant temperature assumption
Isotopic composition of the ocean today: δ18O = 0 ‰
Assume ∆δ18Oglacial-interglacial = 1.7 ‰, and the average isotopic
composition of the ice sheets, δ18Oice = -35 ‰
Let y = ∆δ18Oglacial-interglacialand x = Vice/Voc
Voc * 0‰ = (V oc – Vice) * y + Vice * (–35‰)
x = y/(y+35‰) = 1.7‰/(1.7‰+35‰) = 0.046
0.046 * 3800 m = 176 m (not this large)
Leads to an increase in average ocean salinity by 1.046 x 35‰ = 36.6‰
Assumes linear relationship between δ18O and ice volume,
and constant mean ice δ18O over glacial-interglacial cycles.
3
Mass balance calculation with a
change in deep ocean temperature
T dependence of equilibrium
isotope fractionation factor
between calcite and water
(0.22‰/ºC)
Contours of ‰ enrichment of
seawater as f(sea-level/icevolume and δ18Oice)
Leads to minimum
δ18Os w = 1.1‰ for a
southern ocean site
Freezing point of deep
seawater ~-1.7ºC
Broecker, 1975
Seawater δ18OG – δ18OI = 1.1-1.7‰
∆TI-G = 0 – 2.5ºC
Sea-level/Ice volume Proxy: δ 18O of
Leg 86 of DSDP in the
Pore Waters
north-west Pacific
http://toxics.usgs .gov /definitions/pore_water.html
McDuff, 1985 Recognized that the
signal is caused by glacialinterglacial cycles dominated by
diffusion (not stratigraphic)
The δ18O of seawater depends on
global ice volume and the mean
isotopic composition of the ice.
McDuff, 1985
4
δ18O of seawater during the LGM
1-D diffusion-advection equation:
∂(δ O)
=
∂t
18
Best fit
∂(δ18O)
18
∂z − ∂(φU f δ O)
φ∂z
φ∂z
∂(φDeff
φ = porosity, Deff =
effective diffusion
coefficient (varies with T
and φ), Uf = advective
velocity
Using McDuff [1985]
data
Unlikely
(high)
advection
rate
Schrag and DePaolo, 1993
Pore fluid constraints on T and δ 18O of
the glacial ocean
Benthic Foraminifera ∆δ18Oglacial-interglacial
∆δ18Oglacial-interglacial ˜ 1.7 ‰
∆δ18Osw = 1‰
1.7‰ – 1‰ = 0.7‰
0.7‰ x 4.2 = ~2-3ºC = ∆Tglacial-interglacial (deep water)
5
Advantages/Disadvantages of Pore
Water δ 18O
Advantages:
Disadvantages:
“Direct” measurement of δ18O
of deep oceanic water
Requires assumptions about
diffusion and advection rates,
and the mean δ18O of
continental ice
Relatively easy measurement
Provides better glacialinterglacial ∆T estimates of
oceanic deep water from
benthic foraminifera
Low resolution (sufficient for
last glacial-interglacial cycle)
Must account for some degree
of spatial variability (one core
does not necessarily give
global average)
δ18O of air O2 from ice cores
δ 18O of air O2 and the Dole Effect
The difference between the two is
called the “Dole effect”
δ18 O2(sea-water) = 0‰
δ18 O2(air) = 23.5‰
Dole effect = δ18O2(air) – δ18O2(sw)
δ18O2(air) is determined by:
1.
2.
Marine photosynthesis (δ18OH2O(sea-water))
Terrestrial photosynthesis (δ18OH2O(leaf-water))
3.
4.
Respiration
Photochemical processes in the stratosphere
Turnover time (e-folding time) of atmospheric O2 with respect to
photosynthesis and respiration is ~1.2 kyr.
6
δ18O of air O2 from ice cores
Bender et al., 1994
June insolation at 20N
(Most variability in
insolation at low latitudes
is associated with
precession, while at high
latitudes tilt also becomes
important.)
∆Dole effect(t) = ∆δ18O2(air)(t) – ∆δ18O2(sw)(t)
∆Dole effect between glacial and
interglacial times = 0‰
“Strong precession
signal (~23 kyr) in
∆Dole effect suggests
that variations in the
Dole effect are
probably linked to low
latitude processes.”
δ18O of air O2 from ice cores
Extracting the ice volume component
Vostok air δ18O record
Benthic foram δ18O record
Tuning targets:
linear
combination of
precession and
obliquity
components in
the proportion
and phases
determined by
cross spectral
“etp” analysis
Shackleton, 2000
7
δ18O of air O2 from ice cores
Extracting the ice volume component
Shackleton, 2000
• Residual obtained by subtracting linearly forced component (41- and
21-kyr variance) from the observed record.
• Difference between residuals gives the ice volume component of the
benthic δ18O record (~1 ‰)
Advantages and disadvantages of air
δ18O2 proxy for ice volume
Advantages
• Record inherently contains
ice volume δ 18O signal
• Globally synchronous due
to long turnover time (~1
kyr) compared to the
mixing time of the
atmosphere (~1 yr)
Disadvantages
• Complications from the Dole
effect, which is not wellunderstood (includes variations
in precipitation δ 18O, balance
between marine and terrestrial
photosynthesis, balance
between photosynthesis and
respiration, changes in
stratospheric O3 chemistry
(such as changes in overhead
column O3 abundance))
8