Deep Ocean Circulation and Climate: Formation of Deep Water Masses and Isotope Signatures | Study notes Geochemistry

Geol. 656 Isotope Geochemistry

Chapter 10

294 8/22/12

the Milankovitch signal, which is the primary

driving force of Quaternary climate oscilla-

tions. We observed that the Milankovitch

variations change only the distribution of so-

lar energy received by the Earth, not the total

amount. If this were the only factor in cli-

mate change, we would expect that the glaci-

ation in the southern and northern hemi-

spheres would be exactly out of phase. This,

however, is not the case. Thus there must be

feedback mechanisms at work capable of

producing globally synchronous climate

variation. Broecker, (1984 and subsequent

papers) argued that one of these was the

deep circulation of the ocean.

The role of surface ocean in climate is well

understood: for example, the south-flowing

California Current keeps the West Coast of

the U.S. relatively dry and maintains more

moderate temperatures in coastal regions

than they would otherwise be. The role of

the deep, or thermohaline, circulation of the

oceans is less obvious, but perhaps no less

important. Whereas the surface ocean circu-

lation is wind-driven, the deep circulation is

driven by density, which is in turn controlled

by temperature and salinity.

In the present ocean, most deep ocean wa-

ter masses “form” in high latitudes. Once

these deep-water masses form, they do not

return to the surface for nearly a thousand years. The principal site of deep-water formation today is

the Southern Ocean where the Antarctic Intermediate Water (AAIW) is formed in the Antarctic Con-

vergence and Antarctic Bottom Water (AABW), the densest of ocean water masses, is formed near the

coast of Antarctica, particularly in the Weddell Sea. A lesser amount of deep water is also formed in

the Labrador, Greenland, and Norwegian Seas of the far northern Atlantic when warm, salty water

from the Gulf of Mexico and the Mediterranean is strongly cooled during winter; this water mass is

called North Atlantic Deep Water (NADW). After formation, this water sinks to the bottom of the ocean

an flows southward. Today, it is the deepest and densest water mass in the North Atlantic. In the

South Atlantic, the somewhat cold and denser AABW flows northward beneath the NADW, which is in

turn overlain by AAIW.

Formation of deep water thus usually involves loss of thermal energy by the ocean to the atmosphere.

Therefore, the present thermohaline circulation of the oceans keeps high latitude climates milder than

they would otherwise be. In particular, energy extracted from the Atlantic Ocean water in the forma-

tion of NADW keeps the European climate relatively mild.

We saw in Chapter 8 that δ13C is lower in deep water than in surface water (Figure 8.18). This results

from biological cycling: photosynthesis in the surface waters discriminates against 13C, leaving the dis-

solved inorganic carbon of surface waters with high δ13C, while oxidation of falling organic particles

rich in 12C lowers δ13C in deep water: in effect, 12C is “pumped” from surface to deep water more effi-

ciently than 13C. δ13C values in the deep water are not uniform, varying with the “age” of the deep wa-

Figure 10.19. Variation in δ18O and δ13C in benthic fo-

raminfera from core RC13-229 from the eastern South

Atlantic. δ13C data suggest the proportion of NADW

in this region increased as the climate warmed. Data

from Oppo and Fairbanks (1987).

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Chapter 10

the Milankovitch signal, which is the primary driving force of Quaternary climate oscilla- tions. We observed that the Milankovitch variations change only the distribution of so- lar energy received by the Earth, not the total amount. If this were the only factor in cli- mate change, we would expect that the glaci- ation in the southern and northern hemi- spheres would be exactly out of phase. This, however, is not the case. Thus there must be feedback mechanisms at work capable of producing globally synchronous climate variation. Broecker, (1984 and subsequent papers) argued that one of these was the deep circulation of the ocean. The role of surface ocean in climate is well understood: for example, the south-flowing California Current keeps the West Coast of the U.S. relatively dry and maintains more moderate temperatures in coastal regions than they would otherwise be. The role of the deep, or thermohaline, circulation of the oceans is less obvious, but perhaps no less important. Whereas the surface ocean circu- lation is wind-driven, the deep circulation is driven by density, which is in turn controlled by temperature and salinity. In the present ocean, most deep ocean wa- ter masses “form” in high latitudes. Once these deep-water masses form, they do not return to the surface for nearly a thousand years. The principal site of deep-water formation today is the Southern Ocean where the Antarctic Intermediate Water (AAIW) is formed in the Antarctic Con- vergence and Antarctic Bottom Water (AABW), the densest of ocean water masses, is formed near the coast of Antarctica, particularly in the Weddell Sea. A lesser amount of deep water is also formed in the Labrador, Greenland, and Norwegian Seas of the far northern Atlantic when warm, salty water from the Gulf of Mexico and the Mediterranean is strongly cooled during winter; this water mass is called North Atlantic Deep Water (NADW). After formation, this water sinks to the bottom of the ocean an flows southward. Today, it is the deepest and densest water mass in the North Atlantic. In the South Atlantic, the somewhat cold and denser AABW flows northward beneath the NADW, which is in turn overlain by AAIW. Formation of deep water thus usually involves loss of thermal energy by the ocean to the atmosphere. Therefore, the present thermohaline circulation of the oceans keeps high latitude climates milder than they would otherwise be. In particular, energy extracted from the Atlantic Ocean water in the forma- tion of NADW keeps the European climate relatively mild. We saw in Chapter 8 that δ^13 C is lower in deep water than in surface water (Figure 8.18). This results from biological cycling: photosynthesis in the surface waters discriminates against 13 C, leaving the dis- solved inorganic carbon of surface waters with high δ^13 C, while oxidation of falling organic particles rich in 12 C lowers δ^13 C in deep water: in effect, 12 C is “pumped” from surface to deep water more effi- ciently than 13 C. δ^13 C values in the deep water are not uniform, varying with the “age” of the deep wa-

Figure 10.1 9. Variation in δ^18 O and δ^13 C in benthic fo- raminfera from core RC13-229 from the eastern South Atlantic. δ^13 C data suggest the proportion of NADW in this region increased as the climate warmed. Data from Oppo and Fairbanks (1987).

Chapter 10

ter: the longer the time since the water was at the surface, the more en- riched it becomes in 12 C and the lower the δ^13 C. Since this is also true of total inorganic carbon and nutrients such as PO 4 and NO 3 , δ^13 C corre- lates negatively with nu- trient and ΣCO 2 concen- trations. NADW has high δ^13 C because it con- tains water that was re- cently at the surface (and hence depleted in 12 C by photosynthesis). Deep water is not formed in ei- ther the Pacific or the In- dian Oceans; all deep waters in those oceans flow in from the South- ern Ocean. Hence deep water in the Pacific, be- ing rather “old” has low δ^13 C. AABW is a mixture of young NADW, which therefore has comparably high δ^13 C, and recircu- lated Pacific deep water and hence has lower δ^13 C than NADW. Thus these water masses can be dis- tinguished on the basis of δ^13 C. Examining δ^13 C in ben- thic foraminifera in cores from a variety of loca- tions, Oppo and Fair- banks (1987) concluded that production of NADW was lower during the last glacial maximum and in- creased to present levels in the interval between 15000 and 5000 years ago. Figure 10.19 shows an ex- ample of data from core RC13-229, located in the South Atlantic. δ^13 C values decrease as δ^18 O increases. As we saw in the previous sections, δ^18 O in marine carbonates is a measure of glacial ice volume and climate. As the climate warmed at the end of the last glacial interval, δ^13 C values in bottom water in the South Atlantic increased, reflecting an increase in the proportion of NADW relative to AABW in this region. From δ^13 C variations in Mediterranean and Central Atlantic cores, Oppo and Fairbanks (1987) also concluded that the production of MIW was greater during the last glacial maximum. Thus the mode of ocean circulation apparently changes between glacial and interglacial times; this change may well amplify the Milankovitch signal.

Figure 10. 20. Cross-section of δ^13 C the North Atlantic today and during glacial times. As discussed in the text, different water masses have dif- ferent δ^13 C signatures. From Curry and Oppo (2005).

Chapter 10 Spring 2011

tween 2° and – 2° C. Be- fore the Eocene, deep water appears to have been much warmer, and thermohaline circulation may have been domi- nated by salinity differ- ences. (The formation of Mediterranean In- termediate Water, which forms as a result of evaporative increase in salinity, can be viewed as a remnant of this sa- linity-dominated circula- tion.) It was probably not until late Miocene that the present thermo- haline circulation was completely established. Even subsequent to that time, important varia- tions may have occurred, as we have seen. The mid-Miocene in- crease in δ^18 O probably represents the expansion of the Antarctic ice sheets to cover West Antarctica. This inter- pretation is supported by δD analyses of sedi- ment pore water. Even though pore water ex- changes with sediment, water dominates the deuterium budget so that δD values are ap- proximately conserva- tive (diffusion also af- fects δD, but this effect can be corrected for). An increase of about 10‰ δD occurs between mid and late Miocene, which is though to reflect the accumulation of deuterium-depleted water in Antarctic ice sheets.

Figure 10.21. Variation in δ^18 O in benthic foraminifera in the Tertiary. Data are from Ocean Drilling Program Sites 659 in the eastern equatorial Atlantic, 588 in the southwest Pacific, 929 in the western equatorial Atlan- tic , 522 in the South Atlantic, and 689 in the Southern Ocean Note that δ^18 O is relative to PDB, rather than SMOW. This is conventional for car- bonates. From Zachos et al. (2001).

Chapter 10 Spring 2011

Continental Isotopic Records

Climate change has left an isotopic record on the continents as well as in the deep sea. As with the deep-sea records, it is the isotopic composition of H 2 O that is the paleoclimatic indicator. The record may be left directly in ice, in carbonate precipitated from water, or in clays equilibrated with water. We will consider examples of all of these in this lecture. As we noted with the deep-sea carbonate record, the preserved isotopic signal can be a function of several variables. Continental records tend to be even more difficult to interpret than marine ones. All the isotopic records we will consider record in some fashion the isotopic composition of precipitation in a given region. The isotopic composition of precipitation depends on a host of factors: (1) The isotopic composition of the oceans (the ice volume effect). (2) The isotopic composition of water in the source area (the δ^18 O of surface water in the ocean varies by a per mil or more because of evaporation, precipitation and freezing and is correlated with sa- linity). (3) Temperature and isotopic fractionation in the source area (when water evaporates a temperature dependence isotopic fraction occurs; kinetic affects will also occur, and will depend on the vigor of mixing of water at the sea surface; higher wind speeds and more turbulent mixing will reduce the kinetic fractionation). (4) Atmospheric and oceanic circulation patterns (as we saw in earlier lectures, the isotopic composi- tion of water vapor is a function of the fraction of vapor remaining, which is not necessarily a sim- ple function of temperature; changes in atmospheric and oceanic circulation may also result in changes in the source of precipitation in a given region). (5) Temperature in the area where the precipitation falls, as this determines the fractionation between vapor and water. (6) Seasonal temperature and precipitation patterns. The isotopic record might reflect water falling during only part of the year, and the temperature recorded may therefore be that of only a single season rather than an annual average. For example, even in a wet area such as Ithaca, recharge of ground water occurs almost entirely in winter; during summer, evaporation usually exceeds pre- cipitation. (7) Evaporation of water or sublimation of ice. The isotopic record might be that of water remaining after some has evaporated. Since evaporation involves isotopic fractionation, the preserved iso- topic record will not necessarily be that of the precipitation that falls. All of these are climatic factors and are subject to change between glacial and interglacial periods. Changes in these factors do not mean that the stable isotope record in a given region is not recording climatic changes, but they do mean that the climatic changes recorded might not be global ones.

Vostok and EPICA Antarctic Ice Cores

Climatologists recognized early on that continental ice preserves a stratigraphic record of climate change. Some of the first ice cores recovered for the purpose of examining the climatic record and ana- lyzed for stable isotopes were taken from Greenland in the 1960’s (e.g., Camp Century Ice Core). Sub- sequent cores have been taken from Greenland, Antarctica, and various alpine glaciers. The alpine gla- ciers generally give isotopic records of only a few thousand years, but are nevertheless useful, re- cording events such as the Little Ice Age. The Greenland and Antarctic cores provide a much longer re- cord. Very long ice cores that covered 150,000 years were recovered by the Russians from the Vostok station in Antarctica in the 1980’s (e.g., Jouzel, et al., 1987) and were deepened over the next 20 years, eventually reaching back 400,000 years. Drilling was halted in 2003 out of concern for intersecting and contaminating the body of water beneath the ice, known as Lake Vostok. Attention then shifted to the EPICA (European Project for Ice Coring in Antarctica) project, which recovered 3270 meters of ice core, extending back through 8 glacial cycles and 8 00,000 years. Hydrogen isotopes show a much larger range and much greater temperature-dependent fractionation that oxygen, so in ice, interest centers of

Deep Ocean Circulation and Climate: Formation of Deep Water Masses and Isotope Signatures, Study notes of Geochemistry

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Chapter 10

Chapter 10

Chapter 10 Spring 2011

Chapter 10 Spring 2011

Continental Isotopic Records

Vostok and EPICA Antarctic Ice Cores