Stable Isotope TracersStable Isotope Tracers Reading: Emerson and Hedges, Chapter 5, p.134-153 OCN 623 – Chemical Oceanography Stable Isotope Tracers •Trace source/sink and pathways

Stable Isotope Tracers

Reading: Emerson and Hedges, Chapter 5, p.134-153

OCN 623 – Chemical Oceanography

Stable Isotope Tracers

• Trace source/sink and pathways of nutrients and chemicals in the ocean.

• Tracers of biological, physical, geological ocean processes

• Record past changes in physical, chemical, and biological processes in the ocean.

• Stable Isotopes - Introduction & Notation

• Isotope Fractionation

• Some Oceanographic Applications

Outline

Isotopes of Elements

The chemical characteristic of an element is determined by the number of protons in its nucleus.

Atomic Number (Z) = number of protons = defines the chemistry

Atomic Mass = protons + neutrons (N)

Isotopes = Atoms with same Z but different N

The chart of the nuclides (protons versus neutrons) for elements 1 (Hydrogen) through 12 (Magnesium)

Valley of Stability

Most elements have more than one stable isotope.

1:1 line

Full Chart of the Nuclides

1:1 lineWith Z>20, number of neutrons becomes greater than the number of protons because Coulomb energy (i.e., electrostatic repulsion) has Z2 dependence

Examples: H, He, C, N and O

All Isotopes of a given element have the same chemical properties, yet there are small differences due to the fact that heavier isotopes typically form stronger bonds and diffuse slightly slower

% Abundance is for the average Earth’s crust, ocean and atmosphere

Element Symbol Protons Neutrons % Abundance Half-life

Hydrogen H 1 0 99.985

D (2H) 1 1 0.015

T (3H) 1 2 10-14 to 10-16 τ1/2 = 12.33 y

Helium 3He 2 1 0.000137

4He 2 2 99.999863

Carbon 12C 6 6 98.89

13C 6 7 1.11

14C 6 8 10-10 τ1/2 = 5730 y

Nitrogen 14N 7 7 99.634

15N 7 8 0.366

Oxygen 16O 8 8 99.757

17O 8 9 0.038

18O 8 10 0.205

Measurement reporting convention (δ or “delta” units):

Generally reported as ratio of a heavy (i.e., rare) isotope to a light (i.e., abundant) isotope

Using a mass spectrometer, isotope ratios can be measured much more precisely than absolute abundances of isotopes

If δ > 0, the sample is enriched in the heavy isotope relative to a standard; If δ < 0, the sample is depleted in the heavy isotope relative to a standard.

If (18O/16O)sample = 0.995 (18O/16O)Standard

δ18O = -0.005 x 1000 = -5‰

€

δ18O =

18O 16O( )Sample

− 18O 16O( )Standard

18O 16O( )Standard

⎡

⎣

⎢ ⎢

⎤

⎦

⎥ ⎥ ×1000 =

18O 16O( )Sample

18O 16O( )Standard

−1

⎡

⎣

⎢ ⎢

⎤

⎦

⎥ ⎥ ×1000

€

δ18O =0.995 18O 16O( )

Standard18O 16O( )

Standard

−1⎡

⎣

⎢ ⎢

⎤

⎦

⎥ ⎥ ×1000

Delta Notation for 18O:

Example:

Element δ value Ratio Standard

Hydrogen δD 2H/1HStandard Mean Ocean Water (SMOW)

Hydrogen δD 2H/1HStandard Light Antarctic Precipitation (SLAP)

Helium δ3He 3He/4He Atmospheric He

Boron δ11B 11B/10B NIST SRM 951

Carbon δ13C 13C/12C Pee Dee Belemnite (PDB)

Nitrogen δ15N 15N/14N Atmospheric N2

Oxygen δ18O 18O/16O

Standard Mean Ocean Water (SMOW)

Oxygen δ18O 18O/16O Standard Light Antarctic Precipitation (SLAP)Oxygen δ18O 18O/16O

Pee Dee Belemnite (PDB)

δ17O 17O/16O Standard Mean Ocean Water (SMOW)

Sulfur δ34S 34S/32S Canyon Diablo Troilite (CDT)

Each isotopic measurement is reported relative to a standard

http://upload.wikimedia.org/wikipedia/commons/e/e4/Canyon-diablo-meteorite.jpg

http://upload.wikimedia.org/wikipedia/commons/c/cf/Meteor.jpg

Isotope Fractionation

• Change in an isotopic ratio that arises as a result of a chemical or physical process

• Types of isotopic fractionation that cause changes in isotopic compositions:

• Equilibrium fractionation• Kinetic fractionation• Mass-independent fractionation

• Fractionation Factor (α):αA-B = RA / RB, where RA and RB are isotope ratios in materials A and B

Example: equilibrium fractionation of oxygen isotopes in liquid water (l) relative to water vapor (g).

H216O(l) + H218O(g) ↔ H218O(l) + H216O(g)

At 20ºC, the equilibrium fractionation factor is:

α = (18O/16O)l / (18O/16O)g = 1.0098

Isotope Fractionation

Equilibrium fractionation

• Exchange reactions in which isotopes are exchanged between two or more species (with isotopic preference).

• Bidirectional (reversible) chemical reactions

• Usually applies to inorganic species

• Temperature dependent

• Generally, the heavy isotope will be concentrated in the phase in which it is most strongly bound (i.e., its lowest energy state).

– Solid > liquid > vapor– Covalent > ionic, etc.

The carbonate system involving gaseous CO2(g), aqueous CO2 (aq), aqueous bicarbonate HCO3- and carbonate CO32-.

It is an important system that exhibits equilibrium isotope effects for both carbon and oxygen isotopes. For example:

13CO2(g) + H12CO3- ↔ 12CO2(g) + H13CO3-

The heavier isotope (13C) is preferentially concentrated in the chemical compound in which it is most strongly bound. In this case 13C will be concentrated in HCO3- as opposed to CO2(g).

For this reaction:

α = (H13CO3- / H12CO3-) / (13CO2 / 12CO2)

α = 1.0092 at 0ºC and 1.0068 at 30ºC

Equilibrium fractionation - Example

176 Chapter 3. Stable Isotope Fractionation

0 5 10 15 20 25 30 35!14

!12

!10

!8

!6

!4

!2

0

2

CO2(g)

CO2(aq)

CO3

2! (Mook, 1986)

CO3

2! (Zhang et al., 1995)

→

→

HCO3

!

T (°C)

ε (o

/oo

)

Figure 3.2.12: Carbon isotope fractionation between the species of the carbonate systemas a function of temperature with respect to HCO−3 . Values according to Mook (1986)and Zhang et al. (1995) are indicated by solid and dashed lines, respectively.

with T being the absolute temperature in Kelvin. The values of these frac-tionation factors as a function of temperature are displayed in Figure 3.2.12(solid lines). It is interesting to note that CO2(g) is enriched in the heavyisotope 13C relative to CO2(aq) as one might expect that the gaseous phasewould be depleted in the heavy isotope relative to the dissolved phase.

The values of εgb and εdg have been measured by different authors,yielding quite similar results (εgb: Mook, 1974; Leśniak and Sakai (1989);Zhang et al., 1995; εdg: Vogel et al., 1970; Zhang et al., 1995; Szaran,1998). For instance, the values and the temperature dependence of thefractionation factors given by Zhang et al. (1995):

εbg := ε(HCO−3 −CO2(g)) = −0.1141 Tc + 10.78%0 (3.2.27)εdg = ε(CO2−CO2(g)) = +0.0049 Tc − 1.31%0 , (3.2.28)

(where Tc is the temperature in ◦C) are very similar to those given by

ε = (α−1) × 103

Zeebe and Wolf-Gladrow, 2001

Kinetic fractionation

• One isotope reacts, diffuses, or evaporates faster than the other.

• Can be due to chemical, physical, or biological processes.

• Occurs during irreversible reactions like photosynthesis, when the rate of chemical reaction is sensitive to atomic mass.

• Essentially all isotopic effects involved with formation / destruction of organic matter are kinetic.

• Usually, the lighter isotope reacts or diffuses faster.

• Reaction products are enriched in the lighter isotopes; reservoir of reactants is depleted in the lighter isotopes.

• Magnitude of isotope effect is temperature and pressure dependent.

12C16O2 (mass = 12 + 2 x 16 = 44) 13C16O2 (mass = 13 + 2 x 16 = 45)

€

Ek =12

mv2

All isotope effects involving organic matter are kinetic. Example:

12CO2 + H2O → 12CH2O + O2 faster13CO2 + H2O → 13CH2O + O2 slower

Thus organic matter gets depleted in 13C during photosynthesis (i.e., δ13C becomes more negative).

Terrestrial plants: ca. -19‰ (range -26 to -7‰)Marine plants: ca. -14‰ (range -22 to -8‰)

so 12CO2 travels 1.1% faster than 13CO2.

These must have the same kinetic energy

€

Ek =12

mAvA2

=12

mBvB2

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vAvB

=mBmA

⎛

⎝ ⎜

⎞

⎠ ⎟

1 2

=4544

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⎝ ⎜

⎞

⎠ ⎟

1 2

= 1.011

Kinetic fractionation - An Ideal Example

• Fractionation occurs when water molecules evaporate from sea surface.

• Equilibrium effect when water molecules condense from vapor to liquid form (rain is heavier than vapor).

• Vapor becomes progressively lighter (i.e., δD and δ18O get lower) with distance from source.

• Evaporation from ocean creates depleted clouds.

• Air masses transported to higher latitude/altitude where it is cooler.

• Water lost due to rain; raindrops are enriched in 18O relative to cloud.

• Cloud gets lighter All values are δ 18O‰

Continent

20 °C

10 °C

Rain0‰

0‰Ocean

–9‰Vapor

–9‰

Rain–8‰

–17‰

Figure 5:11: Schematic diagram

of the oxygen isotopic fractionation

among seawater, the atmosphere

and rain on land. See figure 5.10

also. Modified from Siegenthaler

(1976).

Emerson & Hedges, 2008

Rayleigh Distillation

Example: Evaporation – Condensation Processesδ18O in cloud vapor and condensate (rain)plotted versus the fraction of remaining vaporfor a Raleigh process. The isotopic compositionof the residual vapor is a function of thefractionation factor between vapor and waterdroplets. The drops are enriched in 18O. The vaporis progressively depleted.

Fractionation increases withdecreasing temperature

Cloud temperature (°C)

Fraction of remaining water

20 15 10 0 –20

1.0 0.8 0.6 0.4 0.2 0

Condensate(rain/snow)

Vapor(cloud)

–25

–30

–20

0

–10

–5

–15

9‰

11‰

δ18 O

(‰

)

Figure 5:10: An idealized

illustration of the differences

between the d18O of condensateand vapor as a function of the

fraction of the remaining water

during the Rayleigh Distillation

process. Envision a cloud that forms

at 20 8C and remains a closedsystem except for water that rains

out as it cools from 20 8Cto! 20 8C. The equilibriumfractionation factor is temperature

dependent, 9% at 20 8C and 11% at0 8C. Modified from Dansgaard(1965).

Rayleigh Distillation

€

RVRV0

=ƒ(α-1)

€

RV is the isotopic ratio of the vapor

RV0 is the initial isotopic ratio of the vapor

ƒ is the fraction of vapor remaining

α is the fractionation factor

Emerson & Hedges, 2008

δ18O in Average Rain vs. Temperature

• 3He to study deep ocean circulation in the Pacific

• 18O to determine freshwater balance in the Arctic Ocean

• 18O as an indicator of the ice ages

Examples: Stable Isotope Applications in Oceanography





3He Plume from East Pacific Rise

Broecker and Peng, 1982

3He at 2500 m depth

located at almost exactly the same water-column depth as the plume maximum, and(ii) the plume helium signal is strongestnear Hawaii. Additional evidence concern-ing the origin of the far-field plume is pro-vided by near-field hydrographic profilescollected directly over Loihi Seamount.The 1985 expedition by the R.V. HakuhoMaru detected hydrothermal plumes overLoihi that were highly enriched in meth-ane, Fe, Mn, Ni, Co, and He (12, 13). The1985 Loihi plume had a two-layer structurewith maxima at depths of -1000 and 1100m, possibly corresponding to separate injec-tion from the shallow Pele's vents versusinjection from deeper hydrothermal vents.In 1992, hydrographic work was conductedby the R.V. T. G. Thompson over LoihiSeamount as well as on the southeast flankof Hawaii and on the Puna Ridge (Fig. 1).A composite plot of 3He/4He values versusdepth for all of these hydrocasts from expe-dition TT-014 (Fig. 3) shows 3He/4He val-ues near background for southeast flank andPuna Ridge stations, whereas the samples

E

100-_

-1000-

-2000-

-3000-

-4000-

-5000-

8( 3He)%15 20 25 30~~~~~-

4'§. Discoverer S94F-. St.30 (20.4°N, 155' W)

St.29 (21.0°N 155' W)St.28 (22.5°N. 155° W)

+ St.27 (24.2 N, 155 W)X-- St.26 (25.8 N, 155 W)- St.25 (27.5°N, 155' W)

Fig. 2. Profiles of 6(3He)% versus depth for severalstations from expedition S94F. Station (St.) loca-tions are shown in Fig. 1. Note that stations 28,29, and 30 all show a clear maximum in 8(3He) ata depth of 1 100 m, which is attributed to ventingon Loihi Seamount. The deeper maximum at adepth of -2000 m is due to hydrothermal ventingon the JdFR and EPR.

0 200,1

-800-

i-E -1000-

at -1200-

.

N ..@00%

S

-1400-J-

-1600-JIFig. 3. Plot of 8(3He)% versustaken over Loihi Seamount anding expedition TT-01 4 in 1992.shown in the insert in Fig. 1.

collected over Loihi Seamount shigh enrichments in 3He/4He in trange between 1000 and 1250these 1995 observations detected tat essentially the same depth asearlier 1985 work. Two of the 199have very high 3He/4He ratios of470 and 410%.

The 17-014 helium data forrntrend that is due to mixing betwground seawater and the hydrothemember (Fig. 4A). A least-squaresto these data gives a slope of '3F(3.25 ± 0.26) x 10-5 or R/RA =1.9 (la errors), which correspon

A 6-

0)

I. 4-E02ov-0 2-

4.0

B 8.0-

- 7.5-

'O- 7.0-E0I l0

Cl' 6.5-

Loihi heliumSlope3He/4He = 3.25x10-,

0/O/

I \\\//t

.-I

MOR heSlope =

4.5 5.0 5.5

4He (0I8cm3/9)

0

0

00

0 @4o 0 100 Discoverer S94F

*St. 22 to 30, 80010St. 22 to 30, 1 70C+St. 1 to 23, all dE

4.1 4.2 4.3 4.4

4He (1 ~8CM3/g)Fig. 4. (A) 3He concentration versus 4He concen-tration for the same suite of TT-01 4 samplesshown in Fig. 3. Concentration units are in cubic

>/O centimeters per gram at standard temperatureand pressure. A linear least-squares fit (long-

400 600 o4010 6qo dashed line) gave a slope of 3He/4He = (3.25 ±0.26) x 10-5, or RIRA = 23.4 + 1.9 (1 a errors).The correlation coefficient for this fit was r2 =0.854. For comparison, a short-dashed line indi-cates the slope of 3He/4He = 1.0 x 10-5 expect-

* ed for helium introduced along the mid-oceanridges (MOR). (B) 3He concentration versus 4Heconcentration for far-field samples collected alongS94F at stations 22 through 30. Shallow samples(at depths of 800 to 1400 m) are shown as solidcircles; deeper samples (from 1700 m to the bot-tom) at the same stations are shown as opencircles. For reference, all the remaining samples

depth for samples from the S94F expedition in the northeast Pacificin the vicinity dur- are shown as crosses. Linear regression fits toCast locations are these data groups show that the shallow Loihi

plume samples belong to a distinct population.

SCIENCE * VOL. 272 * 17 MAY 1996

how very helium isotope ratio of the pure end-mem-the depth ber fluid. This value is in good agreementm. Thus, with the end-member helium ratio of R/RA:he plume = 23.1 ± 2.3 determined by Kodera et al.s did the (16) from their 1985 samples and also5 samples agrees with direct measurements of helium8(3He) = isotope ratios in Loihi vent fluids (9, 10).

The fact that a near-field plume was ob-a a linear served over Loihi in 1983 ( 11), in 1985 (12,een back- 13), and again in 1995 (this work) indicatesrmal end- that the plume is being maintained by thelinear fit Loihi vents in a steady-state mode at aIe/4He = depth of - 1100 m.= 23.4 ± The elevated 3He/4He ratio (R = 27 RA)ds to the of the helium emanating from Loihi Sea-

mount suggests that it might be possible todetect this elevated isotope ratio in thefar-field plume samples. As shown in Fig.

,/ 4B, the shallow helium plume samples (atdepths of 800 to 1400 m) from S94F aredistinct from the other samples, correspond-ing to a higher enrichment in 3He relativeto 4He. This is exactly what would be ex-pected for a plume originating from an end-member fluid with elevated 3He/4He such

-~ as Loihi. Thus, the far-field plume samplesAum are a diluted version of the near-field sam-1.1 x10-5 ples collected directly over Loihi (Fig. 4A).

Because of the analytical uncertainty in the6.0 absolute concentrations, it is not possible to

accurately determine the end-member 3He/4He ratios by means of linear regression fitsto the data points in Fig. 4B. However,linear regression fits show that the shallowsamples and the deep samples belong to twodistinct populations at better than 70%confidence limits. This detectable isotopicanomaly relative to other Pacific watersamples is a further indication that thisshallow helium plume has its origin onLoihi Seamount.

to 1400 m The fact that the helium signal from)m to bottomepths

Fig. 5. 8(3He)% contoured in section view forWOCE line P17 along 1 35°W. In this section, theLoihi plume appears as an upturning of the 8(3He)contours in the depth range from 900 to 1400 m,between latitudes 150 and 25°N as indicated bythe dashed ellipse. The plume core centered at8°N and a depth of 2500 m is helium from theEPR, whereas the weaker signal north of 32°Nand at a depth of 2000 m is from the JdFR.

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Lupton, 1996

JdFREPR

1000 2000 3000 4000 5000 6000 7000

120°E

120°E

180°!

180°!

120°W

120°W

60°W

60°W

80°S 80°S

60°S 60°S

40°S 40°S

20°S 20°S

0°! 0°!

20°N 20°N

40°N 40°N

60°N 60°N

80°N 80°N

S3

P11

P17

P17

P17

P10 P13P9P8

P14

P14

P16

P18

P19

A21

S4

P1W

P1

P3

P4

P6

P17E

P17NE

P31

P24

P17CA

P21

P17CCA

P13J

P2, P2T

P15P15SA

P15

120°E 180°! 120°W 60°W80°S

60°S

40°S

20°S

0°!

20°N

40°N

60°N

80°S

60°S

40°S

20°S

0°!

20°N

40°N

60°N

80°N 80°N120°E 180°! 120°W 60°W

S3

P11P15P15SA

P17

P17

P17

P10 P13P13JP9

P8

P14

P14

P16

P18

P19

A21

S4

P1W

P1

P2

P3

P4

P6

P17E

P17NE

P31

P24

P17CA

P21

P17CCA

P15

Oceanographic Sections and Stations used in the Pacific WOCE Atlas

Loihi is detectable on the 1100-m depthhorizon some 400 km north of Hawaii sug-gests that the Loihi plume might be muchmore extensive than previously imagined.One indication of the lateral extent of theLoihi plume is the excess 'He above theregional background that is present at adepth of 1100 m on World Ocean Circula-tion Experiment (WOCE) line P17 at135°W over 2000 km to the east of Hawaii(Fig. 5). The Loihi helium plume does notappear as a distinct maximum in these P17profiles as it does in the stations closer toHawaii (Fig. 2). Instead, the Loihi plumehere is expressed as an upward tum of the8(3He) contours at depths of 1000 to 1500m between 14°N and 24°N (Fig. 5). Be-cause of the large distance between theproposed source and these P17 stations, it isdifficult to demonstrate that this signal at135°W is from Loihi. For example, thisshallow 3He could be due to shallow inputfrom a hydrothermally active seamount atanother location. However, the excess 3Heat 135°W is at the correct depth and lati-tude to be attributed to hydrothermal inputon Loihi.A more comprehensive view of the

Loihi plume is provided when helium datafrom several different expeditions are com-bined. Figure 6, which shows 8(3He)% con-toured on a surface at a depth of 1100 m(potential density cu, _ 27.3) (17), includessamples from expeditions CGC-91, S94F,and TPS-24, and WOCE lines P4, P16,P17, P18, and P19C (14). The Loihi 'Hesignal is detectable in several of the TPS-24stations north and east of Hawaii but isabsent in all of the WOCE P4 samplesalong 8.5°N. Thus, it seems that the cur-rents at this depth do not transport theLoihi hydrothermal signal as far south as80N.

The asymmetric 3He distribution shown

in Fig. 6 implies that transport is from westto east at latitude 20°N at depths of 900 to1300 m. This flow is in the opposite direc-tion to that suggested by the deeper 3Heplumes, which trend westward into the in-terior of the Pacific basin from injectionsites along the EPR and JdFR. For example,the deep helium section along 10°N (4)clearly indicates westward transport of 3Hefrom venting sites on the EPR axis at-105°W. This westward flow at 10°N isconsistent with observations of asymmetricplume behavior during near-field investiga-tions on the ridge axis (18). The heliumplume from the JdFR also trends southwestinto the interior of the Pacific basin fromthe injection sites in the far northeast Pa-cific (3). However, these plumes from mid-ocean ridge hydrothermal activity may betracing a much different part of the circu-lation regime because they are much deeper(2000 to 2500 m in depth) than the Loihiplume.

The Loihi helium plume places somestrong constraints on the intermediatedepth circulation of the central north Pa-cific. The near-surface circulation of thenorth Pacific Ocean has been described interms of two gyres: one large, midlatitudeanticyclonic gyre, and another higher lati-tude, cyclonic gyre. Reid (19) notes thatthe centers of these two gyres move fartherpoleward with increasing depth. A map ofsteric height at 500 dbar relative to a 1500-dbar surface (19, 20) shows a flow that iswestward in the vicinity of Hawaii, corre-sponding to the southem limb of this sub-tropical gyre. A similar map of steric heightat 1000 dbar relative to 3000 dbar alsoindicates westward flow in the vicinity ofHawaii (21). These interpretations of thecirculation are clearly at variance with thehelium data. However, Yoshida andKidokoro (22) and Reid and Mantyla (23)

Fig. 6. 8(3He)% contoured on a surface at a depth of 1 100 m, showing the broad lateral extent of the Loihiplume. In some cases, bottle data were interpolated to 1 1 00-m deep surface. The contour interval is 1 %in 8(3He); the accuracy of the measurements is 0.25% (1 cr). This figure includes data from eight differentexpeditions spanning the time interval from 1985 to 1994. Although these data are not synoptic, thesampling period is relatively short compared with the time scale for circulation at this depth. Helium dataalong WOCE lines P4 and P16 were provided by W. Jenkins (4, 23).

discuss an additional eastward flow at adepth of -1000 m centered at -20'N.This so-called subtropical counter-currentagrees quite well with the eastward trans-port suggested by the Loihi helium plumedistribution. Additional helium samplingplanned in the vicinity of the Hawaiianislands will enhance our picture of theLoihi plume, thereby providing furtherconstraints on the existence of thiscounter-current and the pattern of shallowcirculation in the central Pacific basin.

REFERENCES AND NOTES

1. J. E. Lupton and H. Craig, Science 214, 13 (1981).2. H. Stommel and A. B. Arons, Deep-Sea Res. 6, 217

(1960).3. J. Lupton, R. Feely, G. Massoth, Eos (fall suppl.) 70,

1131 (1989); J. E. Lupton, J. Geophys. Res. 95,12829 (1990).

4. Helium and tritium data from WOCE line P4 havebeen reported by W. J. Jenkins in Data Release 6.0(WHOI Helium Isotope Laboratory, 25 August 1995),which is available at http://kopernik.whoi.edu/data/tps.html; some of these data are also reported in S.E. Wijffels et al., Deep-Sea Res., in press.

5. H. Craig, V. Craig, R. Comer, J. Costello, K. Farley,Eos (fall suppl.) 69, 1478 (1988); J.-L. Cheminee etal., Earth Planet. Sci. Lett. 107, 318 (1991).

6. A. Malahoff, in Volcanism in Hawaii, R. W. Decker etal., Eds. (U.S. Government Printing Office, Washing-ton, DC, 1987), p. 133; D. M. Karl, G. M. McMurtry,A. Malahoff, M. 0. Garcia, Nature 335, 532 (1988); P.N. Sedwick, G. M. McMurtry, J. D. Macdougall,Geochim. Cosmochim. Acta 56, 3643 (1992); F. J.Sansone, C. G. Wheat, G. M. McMurtry, J. E. Lup-ton, G. P. Klinkhammer, Eos (fall suppl.) 75, 313(1994).

7. M. D. Kurz, W. J. Jenkins, S. R. Hart, D. A. Clague,Earth Planet. Sci. Lett. 66, 388 (1983); W. Rison andH. Craig, ibid., p. 407; I. Kaneoka, N. Takaoka, C. A.Clague, ibid., p. 427.

8. J. E. Lupton, Annu. Rev. Earth Planet. Sci. 11, 371(1983).

9. H. Craig, J. A. Welhan, D. R. Hilton, Eos (fall suppl.)68, 1553 (1987).

10. P. N. Sedwick, G. M. McMurtry, D. R. Hilton, F. Goff,Geochim. Cosmochim. Acta 58, 1219 (1994); J. E.Lupton, F. J. Sansone, C. G. Wheat, unpublisheddata from 1993 Pisces dives on Loihi Seamount.

11. Y. Horibe, K.-R. Kim, H. Craig, Eos (fall suppl.) 64,724 (1983); K.-R. Kim, H. Craig, Y. Horibe, ibid., p.724.

12. H. Sakai et al., Geochem. J. 21, 11 (1987).13. T. Gamo, J.-l. Ishibashi, H. Sakai, B. Tilbrook,

Geochim. Cosmochim. Acta 51, 2857 (1987).14. The data presented here were collected on nine sep-

arate expeditions as follows: TPS-24, R.V. T. G.Thompson, 1985; WOCE P4, R.V. Moana Wave,1989; WOCE P17, R.V. Thomas Washington, 1991;WOCE P16, R.V. Thomas Washington, 1991; CGC-91, NOAA Ship Discoverer, 1991; TT-01 4, R.V. T G.Thompson, 1992; WOCE P19C, R.V. Knorr, 1993;S94F, NOAA Ship Discoverer, 1994; and WOCEP18, NOAA Ship Discoverer, 1994.

15. With the exception of expedition TT-014, the sam-ples in this study were hermetically sealed into thecopper tubing by cold-welding under high pressurewith a hydraulic press [C. Young and J. E. Lupton,Eos (fall suppl.) 64, 735 (1983)]. This method pro-vides long-term sample storage without loss of in-tegrity. However, for expedition TT-01 4, sampleswere sealed into copper tubing with standard refrig-eration clamps [J. E. Lupton, Earth Planet. Sci. Lett.32, 371 (1976)]. Helium isotope measurements weremade on a dual-collector, statically operated massspectrometer at the Hatfield Marine Science Center,Newport, OR. The estimated precision is 1 a0.25% in 8(3He), and 1 r = 0.5% in 3He and 4Heabsolute concentrations. However, because of leak-

SCIENCE * VOL. 272 * 17 MAY 1996978

on

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3He at 1100 m depth





( )P. Schlosser et al. ! The Science of the Total En"ironment 237!238 1999 15!30 25

18 Ž .Fig. 7. Distribution of " O in the surface waters depth !15 m of the Arctic Ocean and the adjacent seas.

Balances of mass, salt, δ18O, and nutrients allow us to separate the contributions of the individual freshwater sources

Freshwater balance in the Arctic Ocean

Schlosser et al., 1999 Ekwurzel et al., 2001

Freshwater balance in the Arctic Ocean

from P. Schlosser

δ18O

River Runoff

Pacific Water

Sea-ice Meltwater





The δ18O of the CaCO3 is a function of:

1) Temp of seawater that foraminifera are growing in:

• Warmer water → lighter δ18O in CaCO3

2) δ18O of seawater that foraminifera are growing in:• Depends on latitude• Depends on sea level

δ18O in Marine CaCO3

shell-‐water

8

12

16

20

24

28

-‐3-‐2-‐101

Tempe

rature (oC)

δ18O (‰)

O. universa

G. bulloides

G. menardii

G. ruber(pink)

G. sacculifer

N. pachyderma (d.)

Temperature dependence of equilibrium fractionation between 16O and 18O during precipitation of CaCO3

less 18O

more 18Oe.g. Bemis et al., Paleoceanography, 1998

δ18O Paleothermometer

Gridded Surface Seawater δ18O

http://data.giss.nasa.gov/o18data/

Note the higher δ18Osw in the evaporation belts and the lower δ18Osw in the high latitudes, which are dominated by excess precipitation.

Any δ18O-temperature relationship depends primarily on the δ18O of the water from which the carbonate is precipitated.

low δ18O

high δ18O

Changes in ice volume also influence δ18O of the ocean

δ18O in marine carbonates and paleotemperature records Vostok ice coreW. M. White Geochemistry

Chapter 9: Stable Isotopes

387 November 30, 2005

Dating of the Vostok ice core was based only on an ice flow and accumulation model. Nevertheless, the overall pattern observed is in good agreement the marine !18O record, particular from 110,000 years to the present. The record of the last deglaciation is particularly similar to that of the marine !18O re-cord, and even shows evidence of a slight return trend toward glacial conditions from 12 kyr to 11 kyr BP, which corresponds to the well-documented Younger Dryas event of the North Atlantic region. It is also very significant that spectral analysis of the Vostok isotope record shows strong peaks in variance at 41 kyr (the obliquity frequency), and at 25 kyr, which agree with the 23 kyr precessional frequency when the age errors are considered. Thus the Vostok ice core data appear to confirm the importance of Milankovitch climatic forcing. It is interesting and significant that even in this core, taken at 78° S, it is primarily insolation at 65° N that is the controlling influence. There are, however, significant differ-ences between the Vostok record and the marine record. Some of these are probably due to inac-curacies in dating; others may reflect differences in northern and southern hemisphere response to or-bital forcing.

9.6.3 Soils and Paleosols

As we found in Chapter 6, the concentration of CO2 in soils is very much higher than in the atmos-phere, reaching 1% by volume. As a result, soil water can become supersaturated with respect to car-bonates. In soils where evaporation exceeds precipitation, soil carbonates form. The carbonates form

Figure 9.26. The lower curve shows smoothed !18O for marine carbonates (SPECMAP), the mid-dle curve shows !D of ice in the Vostok ice core, and the upper curve the temperature calculated from !D at Vostok at the time of deposition of the ice relative to the present mean annual tempera-ture. Vostok data is from Jouzel, et al., (1987, 1993, 1996).

Warmer/less ice

Colder/more ice

444

Palaeoclimate Chapter 6

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Figure 6.3. Variations of deuterium (!D; black), a proxy for local temperature, and the atmospheric concentrations of the greenhouse gases CO2 (red), CH4 (blue), and nitrous oxide (N2O; green) derived from air trapped within ice cores from Antarctica and from recent atmospheric measurements (Petit et al., 1999; Indermühle et al., 2000; EPICA com-munity members, 2004; Spahni et al., 2005; Siegenthaler et al., 2005a,b). The shading indicates the last interglacial warm periods. Interglacial periods also existed prior to 450 ka, but these were apparently colder than the typical interglacials of the latest Quaternary. The length of the current interglacial is not unusual in the context of the last 650 kyr. The stack of 57 globally distributed benthic !18O marine records (dark grey), a proxy for global ice volume fl uctuations (Lisiecki and Raymo, 2005), is displayed for comparison with the ice core data. Downward trends in the benthic !18O curve refl ect increasing ice volumes on land. Note that the shaded vertical bars are based on the ice core age model (EPICA community members, 2004), and that the marine record is plotted on its original time scale based on tuning to the orbital parameters (Lisiecki and Raymo, 2005). The stars and labels indicate atmospheric concentrations at year 2000.

IPCC Fourth Assessment Report, Working Group I Report "The Physical Science Basis", Chapter 6 Palaeoclimate

ocean overturning and increased surface temperatures should have decreased the flow of dissolved oxygen to deep water. Several direct lines of evidence, such as laminated sediment in cores from the Car-ibbean and central Arctic regions, suggest that dissolved oxygen did indeed decrease across the PETM. Moreover, the PETM coincided with a major extinction of benthic foraminiferans, with widespread oxygen deficiency in the ocean as a possible cause17.

With such ocean conditions, greater preservation and burial of solid organic carbon in deep-sea sediments might be predicted, effectively countering the decreased carbon flux from surface waters. However, this has not been documented. Two largely unexplored processes involving the microbial decomposition of organic carbon, both functioning as additional positive feedbacks, might operate during times of massive carbon input and rapid warming. Carbonate dissolution in the deep ocean decreases sedimentation rates, exposing organic carbon at or near

the sea floor for a longer duration, and warming of deep waters will accelerate overall microbial activity and the consumption of organic carbon. Future investigations might therefore focus specifically on the evidence for changes in ocean overturning, oxygen deficiency and the burial of organic carbon.

The positive feedbacks of greatest concern for understanding overall global warming may be those that could release hundreds to thousands of gigatonnes of carbon after initial warming11–13. The large masses of organic carbon stored in soils (for example, as peat) or sediments of shal-low aquatic systems (for example, wetlands, bogs and swamps) represent a potential carbon input, should regions that were humid become drier. Rapid desiccation or fire could release carbon from these reservoirs at rates faster than carbon uptake by similar environments elsewhere. By contrast, regions that once were dry might emit methane as they become wetter18. Methane might also enter the ocean or atmosphere through the

–1

0

1

2

3

4

5

0 10 20 30 40 50 60

Miocene Oligocene

4

0

8

12Antarctic ice sheets

Full scale and permanentPartial or ephemeral

PETM(ETM1)

4,000

0

1,000

2,000

3,000

5,000

Boron

Alkenones

0a

b

10 20 30 40 50 60

Anthropogenic peak (5,000 Gt C)

ETM2

PalaeoceneEocene

Mid-EoceneClimatic Optimum

Early EoceneClimatic Optimum

NahcoliteTrona

Northern Hemisphere ice sheets?

Plio-cene

Pleistocene

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CO2 proxies

Figure 2 | Evolution of atmospheric CO2 levels and global climate over the past 65 million years. a, Cenozoic pCO2 for the period 0 to 65 million years ago. Data are a compilation of marine (see ref. 5 for original sources) and lacustrine24 proxy records. The dashed horizontal line represents the maximum pCO2 for the Neogene (Miocene to present) and the minimum pCO2 for the early Eocene (1,125 p.p.m.v.), as constrained by calculations of equilibrium with Na–CO3 mineral phases (vertical bars, where the length of the bars indicates the range of pCO2 over which the mineral phases are stable) that are found in Neogene and early Eocene lacustrine deposits24. The vertical distance between the upper and lower coloured lines shows the range of uncertainty for the alkenone and boron proxies. b, The climate for the same period (0 to 65 million years ago). The climate curve is a stacked deep-sea benthic foraminiferal oxygen-isotope curve based on records from

Deep Sea Drilling Project and Ocean Drilling Program sites6, updated with high-resolution records for the interval spanning the middle Eocene to the middle Miocene25–27. Because the temporal and spatial distribution of records used in the stack are uneven, resulting in some biasing, the raw data were smoothed by using a five-point running mean. The δ18O temperature scale, on the right axis, was computed on the assumption of an ice-free ocean; it therefore applies only to the time preceding the onset of large-scale glaciation on Antarctica (about 35 million years ago). The figure clearly shows the 2-million-year-long Early Eocene Climatic Optimum and the more transient Mid-Eocene Climatic Optimum, and the very short-lived early Eocene hyperthermals such as the PETM (also known as Eocene Thermal Maximum 1, ETM1) and Eocene Thermal Maximum 2 (ETM2; also known as ELMO). ‰, parts per thousand.

281

NATURE|Vol 451|17 January 2008 YEAR OF PLANET EARTH FEATURE

Stacked deep-sea benthic foraminiferal oxygen-isotope curve. The δ18O temperature scale, on the right axis, was computed on the assumption of an ice-free ocean; it therefore applies only to the time preceding the onset of large-scale glaciation on Antarctica (about 35 million years ago).

δ18O as Indicator of Climate for the Past 65 Million Years

Zachos et al., 2008

Documents

Stable Isotope TracersStable Isotope Tracers Reading: Emerson and Hedges, Chapter 5, p.134-153 OCN 623 – Chemical Oceanography Stable Isotope Tracers •Trace source/sink and pathways