Stable Isotope Geochemistry II Lecture 31 Mass Dependent

Stable Isotope Geochemistry II Lecture 31

Mass Dependent Fractionation • • • If a 4‰ fractionation of δ 18 O is observed in a particular sample, what value of δ 17 O do we predict? We might guess it would ½ as much. Mass occurs in a variety of ways in the partition function, as m 3/2, as reduced mass, and in the exponential term. Consequently, the ratio of fractionation of 17 O/16 O to that of 18 O/16 O in most cases is about 0. 52. Nevertheless, the fractionation between isotopes predicted by this equation is proportional to the difference in mass – this is referred to as mass-dependent fractionation. There are some exceptions where the ratio of fractionation of 17 O/16 O to that of 18 O/16 O is ≈1. Since the extent of fractionation in these cases seems independent of the mass difference, this is called massindependent fractionation. Examples o o o Oxygen in meteorites Sulfur in Archean sulfides Oxygen in stratospheric gases

Mass Independent Fractionation • • • Most examples of ‘MIF’ seem to be related to photochemical reactions. Formation of ozone in the stratosphere involves the energetic collision of monatomic and molecular oxygen: O + O 2 → O 3 The ozone molecule is in a vibrationally excited state and subject to dissociation if it cannot lose this excess energy. The excess vibrational energy can be lost either by collisions with other molecules, or by partitioning to rotational energy. In the stratosphere, collisions are infrequent, hence repartitioning of vibrational energy represents an important pathway to stability. Because there are more possible energy transitions for asymmetric species such as 16 O 18 O and 16 O 17 O than symmetric ones such as 16 O 16 O, the former can more readily repartition its excess energy and form a stable molecule. In the troposphere, collisions more frequent, reducing this effect.

Isotope Geothermometry • One of the principal uses of stable isotopes is geothermometry. Stable isotope geothermometers are based on the temperature dependence of the fractionation factor or equilibrium constant, which can generally be expressed as: o • (at low temperatures, the form of changes to α ∝ 1/T). Temperature dependence can be theoretically calculated or experimentally measured. Measuring the isotopic composition of two phases allows us to calculate the temperature at which they equilibrated (assuming, of course, that they did equilibrate).

Isotope “Clumping” • Consider the distribution of 18 O between CO and O 2 (Example 9. 1). CO and O 2 will actually consist of 12 isotopically distinct molecules or “isotopologues”, such as 12 C 16 O, 12 C 17 O, 13 C 18 O, 16 O 17 O, etc. • The distribution of isotopes between these species will not be random but rather some of these isotopologues will be thermodynamically favored. o Essentially, grouping the heavy isotopes in one molecule, e. g. , 13 C 18 O, reduces bond energy by a bit more than twice the reduction of putting one heavy isotope in the molecule. Thus “clumping” of heavy isotopes reduces system energy. • This ‘clumping’ depends on temperature (greater at low T). By analyzing the various isotopologues of the species, one can calculate equilibrium temperatures. o The advantage is that we need analyze just one phase involved in the reaction, for example, carbonate precipitated from water. o In addition to calculating temperature, one can also calculate the isotopic composition of the water. • This is a very new field, but holds great promise in isotope geothermometry.

Stable Isotopes in the Hydrosphere and Biosphere

Isotopic Fractionation in the Hydrosphere • δ 18 O in the oceans is nearly constant at 0. However, as a consequence of Rayleigh fractionation and temperature dependence of the water-vapor fractionation factor, the isotopic composition of precipitation varies widely. • There is a strong correlation with temperature.

Isotopic Fractionation in the Hydrosphere • Distance from source and orographic effects also affect δ 18 O of precipitation. • δD is strongly correlated with δ 18 O in precipitation. This is known as the Meteoric Water Line.

Biological Fractionations • • • Relatively large carbon isotope fractionations occur during photosynthesis as a result of: o o More rapid diffusion of 12 CO 2 into the leaf (∆ ≈4. 4‰) and The 12 C-O bond being weaker than the 13 C-O bond, 12 CO 2 is preferentially fixed. Key step in ‘C 3’ plants is reaction of ribulose bis-phosphate with CO 2 to make 2 molecules of a 3 -carbon chain, 3 phosophglyserate (which is subsequently converted to sugar and the phosphate recycled). o The result is a fractionation of -20 - to -30‰. In ‘C 4’ plants, which include grasses, sugar cane and maize, CO 2 is initially fixed into a 4 -carbon chain and then transported into an interior cell where it is converted to 3 phosophglyserate. This process more efficiently utilizes CO 2 and thus results in less isotopic fractionation. o C 4 plants are heavier on average, δ ≈ -13‰ Only small subsequent fractionations occur as sugars are converted into other organic substances and during respiration. As a consequence, the isotopic composition of plants reflects the photosynthetic fractionation and this isotopic composition is maintained up the food chain (with 1 or 2‰ fractionation at each step). The isotopically light signature of photosynthesis is also retained in sedimentary organic matter, including oil and gas (although fractionation during catagensis results in ‘thermogenic’ methane being even lighter).

13 δ C • • • There is a fractionation with dissolution of CO 2 in water so that dissolved CO 2 is heavier. δ 13 C is high in dissolved inorganic carbon (DIC, which is mainly HCO 3–) in surface waters due to preferential photosynthetic utilization of 12 C. δ 13 C is lower in deep water as falling organic matter is ‘remineralized’ (metabolized). o • in the Oceans δ 13 C of deep water decreases with age, so it can be used as a water mass tracer. Incorporation into carbonate shells allows for use in paleooceanography. Variable productivity results in somewhat variable δ 13 C surface water.

Nitrogen & Sulfur Isotope Fractionation • • Biological utilization results in N isotope variations. Many are related to redox reaction between N in 5 different valence states. N incorporated into living matter as amine groups (NH 2–) but can be taken up in either reduced (NH 3) or oxidized (NO 3 -) form - not as N 2 (except by bacteria). Terrestrial plants on the whole has slightly ‘heavy’ nitrogen - but contamination by artificial fertilizer has screwed this up. Extent of fractionation depends on availability. o • Not much fractionation when it is scarce as essentially all is utilized. There are particularly large fractionations, 20 to 40‰, between sulfide and sulfate. These are often biologically mediated.

Isotopic Fingerprints • Plants can be grouped based on their carbon and isotopic compositions due to different fractionations of marine, C 4, and C 3 autotrophs and nitrogen fixers (cyanobacteria and legumes) and nonnitrogen fixers (everything else).

You are what you eat • • • Carbon isotopes undergo 1 -2‰ fractionation with each step up the food chain. Given the large variation in autotrophs, these changes are small. N isotopes also undergo 1 -2‰ fractionation with each step. These are relative large compared to differences in autotrophs. Marine food chains tend to be longer, so isotopic differences at the top are greater.

Isotopes in Archeology • Pioneering work by De. Niro (not that one) and his advisor Epstein (a student of Urey) has led to widespread use of isotopes in archeology and paleontology. • Bone collagen, teeth, and cooking residues on pottery shards retain original δ 13 C and δ 15 N signatures. This allows reconstruction of diets. • 87 Sr/86 Sr and δ 18 O can “fingerprint” the region an individual lived in as these too are preserved in bone.

Appearance of C 4 Plants • C 4 plants (again, mainly warm climate grasses) first became important in the Miocene. o This is documented by, among other things, a shift is δ 13 C in soil carbonates in Pakistan (similar shifts seen in N. America). • Question is why? o C 4 pathway more efficient at low atmospheric CO 2 levels (otherwise not), so it may be an adaptation to lower Neogene CO 2 levels.

Isotopic Fossils • The earliest evidence (still somewhat disputed) of life is low δ 13 C values in graphite inclusions in apatite found in 3. 8 Ga metasediments from Isua, Greenland. • δ 13 C as low as -19‰ occur. There is no demonstrated natural mechanism (so far) to produce such isotopically light carbon abiologically.

Stable Isotopes in Paleoclimatology

Water-Carbonate Fractionation • Urey calculated the temperature dependence of the water-carbonate δ 18 O fractionation and pointed out it could be used as a paleothermometer by solving for T: • T (˚C) = 16. 9 -4. 2∆+0. 13∆2 o where ∆ is the difference between calcite and the water it precipitated from. • He then had his students perform experiments to verify predictions.

Quaternary • • • Urey’s student, Cesar Emilliani, analyzed δ 18 O in forams from a variety of deep-sea cores and reported 15 glacial cycles in the last 600, 000 years in his 1955 dissertation. Subsequent work greatly refined this record, leading to a standard δ 18 O curve in the late 1970’s. Emilliani had noticed the cyclicity in the curves and concluded that Milankovitich’s theory of climate change was correct: it was caused by changes in the Earth’s orbit and rotation. 18 δ O Record

Deducing Temperature Change • Two factors change δ 18 O in marine carbonates: o o • In order to determine temperature changes, one must know how the isotopic composition of water changed. o o • Temperature dependence of the fractionation factor carbonate will be heavier at lower T. Storage of isotopically light water on continents as glaciers. Consequently, seawater, and also carbonates, will be heavier during glacial periods. Deep water temperature changes less, so benthic forams provide some control on this. Ice volumes can be determined from sealevel change (subsequently constrained by dating coral reefs with U-Th). In addition, of course, it is necessary to accurately date strata in the cores. o Has evolved from extrapolating 14 C dates and magnetostratigraphy to more sophisticated approaches like U-Th and 10 Be, etc.

Milankovitch Theory • Earth’s orbit and rotation vary regularly in 3 ways: o The obliquity of the rotational axis relative to the orbital plane. o Eccentricity of the orbit o Precession: the direction the Earth’s rotational axis points at perigee and apogee of orbit. • These factors influence the distribution of solar energy (insolation) in time and space over the course of a year, but do not change global annual insolation. • ‘Milankovitich parameters’ are well determined from astronomical observations (have been known for a very long time).

• • Imbrie, Hayes and others Model Imbrie and colleagues (1976, 1985) applied Fourier analysis to the standardized δ 18 O curve (CLIMAP project) to deduce the primary frequencies (dividing into two parts, <400 ka and >400 ka). They then build a model where each Milankovitch frequency influenced climate with a different phase and gain. The model accounted for r 2 = 0. 77 of the observed variance in δ 18 O. This kind of model has, of course, been greatly subsequently enhanced with better data, GCM’s, ocean circulation models, etc.