Cavity ringdown spectroscopy, instrument

Cavity Ringdown Spectroscopy, instrument

cost ~$90K
vs. ~$400K for mass spec
from Gupta, P. 2009, Chapman Conf. poster


Cavity Ringdown Spectroscopy, principles

from picarro.com

-

idea is to compare empty chamber
and full-chamber ring-down
across several absorption
lines

-

must determine unknowns against
a calibration of ring-downs of known standa


Geothermometry & paleoclimate proxies

10/10/12

The JOIDES Resolution
drillship

Temperature-dependent fractionation - recap
Equilibrium fractionation is temperature-dependent, always.
- we’ve discussed the liquid-vapor fractionation for precipitation
- today: carbonate-liquid fractionation

any solid phase
red=warm
blue=cold;
arrows track movement of 18O through phase changes


Carbonate δ18O – introduction
Minerals (e.g. carbonate, quartz, barite, etc) form from super-saturated solution.
δ18O of these minerals is a fxn of δ18O of solution and temperature of solution
Remember: the δ18O of a solid phase is usually reported in PDB (heavy standard)
while δ18O of liquid phase is usually reported in SMOW (light standard)
interconversion equation:

δ 18OSMOW = 1.03086(δ 18OPDB ) + 30.86

Friedman and O’Neil (1977)

Important: You need to know the δ18O of the solution to derive temperature from δ18Osolid
The ocean δ18O is defined as 0‰ (Standard Mean Ocean Water), and it’s a big volume,
so how do you change δ18O of seawater?


Carbonate δ18O – temperature relationships
The relationship between water-δ18O, temperature, and the equilibrium

δ18O of calcite was determined empirically by Sam Epstein et al., (1953)
and later modified by Craig (1965): T °C = 16.9 − 4.2(δ − δ ) + 0.13(δ − δ ) 2
c
w
c
w
O’Neil et al. (1969) determined an experimental relationship for the
temperature-dependence of α for the calcite-water system:
103 ln α = 2.78(106 T −2 ) − 2.89

NOTE
δ c must be wrt PDB,
δ w must be wrt SMOW
good for low T,
paleoceanography

T in Kelvin
good for high T


Aragonite δ18O – temperature relationships
Why is the δ18Oarag-water α
different than the δ18Ocal-water α?

Is the α larger for aragonite or calcite?
T-dependence of α (T in Kelvin):

Zhou & Zheng, GCA, 2003

LGM

Glacial-Interglacial foraminifera δ18O, revisited
Data from deep-sea
(benthic) foraminifera
show +1.5‰ δ18O
shift during LGM

The million-dollar question in paleoceanography:
How much of this shift was due to ice volume (sea level change) and how much
was due to temperature change?
Schrag modelled the glacial-interglacial shift in porewater δ18O (~1.0‰),
so we have 0.5‰ left over for temperature change.
How much did bottom water temperatures change during the LGM? (problem set)
Or you could measure temperature (trace metal concentrations in carbonates),
and obtain a “residual” δ18O that gives you the δ18OSW change.


Complications: Kinetic effects, vital effects and carbonate δ18O
Fact: very few organisms precipitate carbonate in isotopic equilibrium with the surrounding
water (the vital effect)
environments”,
Kinetic isotope

effects underlie
vital effect

One problem: skeletons are precipitated in super-saturated “microwith sources from surrounding water & metabolic products
Another problem: isotopic exchange may be rate-limited in biological reactions

Can track kinetic effects
with isotope-isotope
plot (δ13C vs. δ18O),
check for slope = 2


values covary with a positive d O/d C slope of 0.29. The foraminiferal d1 8 O/d1 3 C relationships obtained herearevirtually identical,
with slopes ranging between 0.29 and 0.33 (Fig. 3). Other ahermatypic corals, calcareous algae and invertebrates such as cidaroid
18
15,16
urchins show oxygen and carbon isotope covariance , although
the slopes can differ from the experimental range. In symbiont-

probably indicate higher calcification rates. Chamber m
for G. bulloides areinconclusive.
Although we cannot offer a definitive mechanism
isotope:[CO 23 − ] relationships, mass balance calculations sh
the results cannot
be explained
by effect,
a simple redistribu
* Not
a very big
isotopes between the dissolved carbon species19. MConna

Carbonate ion effect on foram δ O
Results from culturing living forams:
increase CO32-, δ18O foram decreases
Constant alkalinity

2−
3

Constant SCO 2

a

4.00

but casts further doubt
on inferring G-I T changes
from
forams
Figure 1 Effect of [CO ] on the d C and d

HL
Dark

4.00

13

18

O

Orbulina universa shell calcite under con

c

HL
Dark

constant alkalinity (a and b) and constant SC

13
Shell d C (‰)

d) conditions. Specimens grown under high li

the dark are shown by open and close
2.00

2.00

the Southern California Bight. Data are gro
0.00

-2.00

0

100

200

300

400

500

600

700

800

-0.50

b

HL
Dark

-1.50

-2.50

-2.50
2–

200

300
2–

400

500

[CO 3 ] (mmol kg–1)

600

200

300

400

500

600

d

the data.
700

800

HL
Dark

Spero et al., Nature 1997

d18O =1.60 — 0.002 [CO3 ] R 2 =0.92

2–

d18O =1.44 — 0.002 [CO3 ] R 2 =0.97

100

100

individual shells. Lines are linear regression

2–

d18O =1.56 — 0.002 [CO3 ] R 2 =0.89

0

2–
d13C =4.29 — 0.006 [CO3 ] R 2 =0.90
2–
d13C =3.30 — 0.006 [CO3 ] R 2 =0.92

-2.00
0

-1.50

-3.50

values 6 1s.d.; most groups are compose

0.00
2–

d13C =4.74 — 0.006 [CO3 ] R 2 =0.95
2–
d13C =3.56 — 0.006 [CO3 ] R 2 =0.93

-0.50

18
Shell d O (‰)

respectively. Arrowheads identify ambient

2–

700

800

-3.50
0

d18O =1.31 — 0.002 [CO3 ] R 2 =0.96

100

200

300

400

500

600

700

800

2–

[CO 3 ] (mmol kg–1)

Figure 2 Effect of [CO23 − ] on the d1 3 C and d1 8 O

Globigerina bulloides chamber calcite under


Glacial-Interglacial climate reconstruction
In order to reconstruct surface temperatures from carbonate δ18O formed
during the LGM, you need to
1) remove the ice-volume effect
2) constrain the δ18O of your local water mass
3) apply the paleo-temperature equation
However, people can use other proxies to get at temperature:
1) foraminifera assemblage data (CLIMAP)
2) tree lines and snow lines will be lower during cold times
3) trace metals in carbonates (Mg/Ca in forams or Sr/Ca in corals)
4) alkenones (saturation index of long-chained alkanes in coccolithophores)

δ18O as tracer of igneous processes
Applications of oxygen isotopes in igneous rocks:
1)
2)

determine temp of formation (water-mineral
or mineral-mineral pairs)
quantify “water-rock” ratios of altered rocks

spectrometer light
intake
A “black smoker”
from the East
Pacific Rise


Oxygen Isotopic compositions of geological materials
Why does eclogite have
a heavier δ18O than MORB?

Lunar rocks
MORB
basic lavas
mantle nodules
eclogites
andesites
Why do metamorphic rocks
ophiolites
exhibit such a range of δ18O?
rhyolites & tuffs

granitic rocks
altered igneous rocks
If carbonates precipitate
metamorphic rocks
from a “light” ocean,
clastic sediments
why are they so heavy?
marine limestones

composition of lunar rocks, carbonaceous
chondrites, and MORB
Why is the ocean so light compared to MORB?

What scale is this?


Principle of Geothermometry in igneous applications
The fractionation of oxygen or hydrogen in different minerals of a rock can be
used as a geothermometer, provided that:
1. minerals deposited at same time, at equilibrium
2. no subsequent alteration
3. fractionation factors and T-dependence known experimentally
NOTE: Using multiple mineral pairs will increase confidence in the calculated temperature,
if the mineral pair temperatures agree – i.e. they are concordant.

For phases m and n:

1000 ln α m − n =

6

a *10
+b
2
T

General form of geothermometry
fractionation equations.

Handy conversions:
1000 ln α m − n = ∆ m − n = δ m − δ n

T in
Kelvin

1000 ln α calcite − water

2.78*106
=
− 2.89
2
T

Remember from
last lecture we talked
about the high-T
water-calcite equation?


T-dependent fractionation in various mineral pairs

Where must these lines converge?

NOTE: These slopes are different – so all you
need to determine T is ∆m-n
The highest fractionation is between
quartz and magnetite.
In general, 18O is increasingly favored
in higher-quartz minerals, and less
favored in hydrous minerals (magnetite,
amphibole, chlorite).

How could we determine the slope of
the Quartz-Muscovite fractionation?


Today’s
Handouts:
Tables of
T-dependent
Fractionation


Example
Quartz, calcite, and chlorite were all precipitated in a hydrothermal vent setting.
Measured δ18O’s:
Quartz: 5.1‰ SMOW
Calcite: 3.8‰ SMOW
Chlorite: -1.5‰ SMOW

Why does quartz have the heaviest δ18O,

and chlorite the lightest?
And why is the quartz only 5.1‰ heavier than SMOW?

Did these minerals precipitate at the same temperature?
How would you begin to solve this problem?


Metamorphism: Water-Rock interactions
Fact: In several places it is possible to measure igneous rocks with δ18O values of -5‰!
These are places were fluid has interacted with the rock (usually at high T) to change
the isotopic composition of the rock.
We can use a mass balance approach to calculate the amount of water that has
reacted with a host rock (or “water/rock ratio”) over time (assuming equilibrium):

for a closed system
What does a “closed
system” mean?

and mass balance equation:

and combining these:

is the equilibrium values for water and mineral,
(need to know temperature independently)

∆ = δw − δr

cwW δ + cr Rδ = cwW δ + cr Rδ r
i
w

i
r

f
w

 cr 
W
δ rf − δ ri
= i
* ÷
f
R δ w − δ r − ∆  cw 

f

cw = conc. of O in water
cr = conc. of O in rock
W = mass water
R= mass rock
superscript I = initial
superscript f = final


Water-Rock interactions II
for an open system
Now we only have a small parcel of water (dW) interacting at any given time, but new
water parcels are injected continuously in time, causing dδr

Rcr d δ r = ( δ wi − [ ∆ + δ r ] ) cw dW
in this scenario we need to integrate to calculate W/R ratios.

 δ rf − δ ri
  cr 
W
= ln  i
+ 1÷ ÷
f
R
 δ w − δ r − ∆   cw 
Probably much more realistic, because water flows through the rock.
In order to solve for W/R interactions, you need to know:
1. the temperature of the interaction (hopefully you can get that by a mineral-mineral pair)
2. the mineral phases that experienced fluid alteration
3. the isotopic composition of the water before it interacted with the rock (δD of rock… why?)
4. the isotopic composition of the rock before it interacted with the water (unaltered samples)


A real-world
example
A characteristic
signature of hydrothermal
activity is a “bulls-eye”
pattern of δ18O values,
with low values in the
middle.
Alteration occurs along
an established conduit
of weakened structures.

Most gems are the product
of low-T, high-fluid
metamorphism –
$1M worth of gold mined
in the Bohemia complex
between 1870 and 1940 happy hunting!


A cool early Earth?
Idea: measure U-Pb dates and δ18O of old zircons
- if you find low δ18O relative to today’s ‘primitive’ mantle, then that implies
interaction with meteoric waters at low temperatures
This work is done using a
laser flourination line plumbed
to a dual inlet mass spec

A Cool Early Earth, (2002) Geology. 30: 351-354.


So what’s causing the relatively high zircon δ18O values at 4.2Ga?
time-line along bottom indicates:
(1) accretion of the Earth,
(2) formation of the Moon and the Earth’s core,
(3) minimum age of liquid water based on high δ18O zircon,
(4) Acasta gneiss, and
(5) Isua metasedimentary rocks.


likely interaction with meteoric waters
which implies a period of few impacts