Carbon Isotopes and Isotopic Fractionation for Complex (e.g., biochemical) Processes

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1 Carbon Isotopes and Isotopic Fractionation for Complex (e.g., biochemical) Processes Reading: White: Chapter 8, the section, "Carbon Isotope Fractionation During Photosynthesis" For more information: Valley and Cole, 2001 Chapters 3 (Hayes), 10 (Des Marais), and 11 (Freeman) Motivation: C, S, N, and other redoxactive elements have massdependent fractionation that is largely controlled by multistep, predominantly biological processes C by photosynthesis S by sulfate reduction N by nitrification, volatilization, denitrification Fe, Cr, Se, and Mo: various redox reactions, may be biologically mediated We need a general framework for understanding how isotopic fractionation varies according to reaction rates and reaction pathways Photosynthesis is the most important driver of C isotope fractionation, but there are other important isotopic effects we must understand in order to use C isotope data Carbon isotopes provide important records for understanding earth history. Guide Questions: In general, how does isotopic fractionation vary multistep chemical reactions that are predominantly forward reactions (far from equilibrium)? How do reaction rates and reaction pathways influence the overall isotopic fractionation of the reaction product relative to the reactant? How does photosynthesis fit this model? What are the various steps involved and how do they differ between C 3 and C 4 plants? What are the resulting ranges of isotopic fractionation? How does CO 2 concentration effect the isotopic fractionation? What are the isotopic fractionations between the dissolved inorganic carbon species, roughly? Are the molecules that make up living tissues in plants and animals all the same in their C isotope composition? If not, how large are the δ 13 C differences between different classes of molecules (roughly)? What are the δ 13 C values of the atmosphere, C 3 plants, C 4 plants, marine organic matter, marine carbonate rock, and methane? The ε value for photosynthesis was apparently lower at times in the past. What does this suggest about conditions at that time? The δ 13 C value of marine carbonate has varied over time? What do larger/smaller δ 13 C values suggest about the marine carbon cycle? How can we constrain animal and human diets in the past using C and N isotopes? In studies of contaminant transport and biodegradation, what two types of information can be gained from C isotope data? Can C and H isotopes be used to determine how a methane deposit formed? If so, what is the general idea?

2 Isotopic fractionation by living cells Right: Highly simplified model of a cell, within which isotopefractionating biochemical reactions occur. Reactant is A and product is B. A could be CO 2, sulfate, nitrate, etc. The various steps have varying isotopic fractionation associated with them; what is the fractionation for the whole process? A B out A inside B inside A+Enzyme Recall, from our work with kinetic isotope effects induced by multistep processes: The ratelimiting step (RLS) is the step with the smallest rate constant; the bottleneck Overall isotopic fractionation (ε) is the sum of all isotopic fractionations up to and including the rate limiting step. Steps after the RLS have little effect on the overall ε Apply this theory to photosynthesis and other biochemical reactions First consider two end member cases: Extreme Case 1: Diffusion into the cell is the ratelimiting step: There is a kinetic effect (small) at the RLS related to diffusion differences. Later steps contribute no fractionation to the overall reaction product flow (flux) Overall fractionation = that of the diffusion step alone. This extreme case could happen only if the enzyme step (A + enzyme => B) is extremely fast so fast that the reactant A is highly depleted inside the cell. Extreme Case 2: Diffusion into the cell is NOT ratelimiting at all. RLS is the enzyme step (A + enzyme => B), which we expect has a large ε becaue bonds are changed. Enzyme step is slow (e.g., if cell is starved and there s little activated enzyme) Diffusion is comparatively fast, and the concentration inside the cell that outside. This strong diffusive connection between the inside of the cell and the pool of reactant, A, outside means that any isotopic fractionation occurring inside the cell will affect the outside reactant pool Thus, the greater reaction rate for lighter isotopes at the enzyme step results heavy isotope depletion in the reaction product and a complementary heavy isotope enrichment in the reactant. This extreme case could happen only when cells are metabolizing very slowly so that diffusion is, in a relative sense, very fast. Intermediate cases: It s just what you would guess. In intermediate cases, the isotopic fractionation is intermediate between the extreme cases. For example, diffusion is often partially ratelimiting, but not completely. In other words, the reaction inside the cell is fast enough so that the rate of influx of reactant via diffusion influences the overall reaction rate and the overall isotopic fractionation but does not control it completely. The result is an ε value somewhere between those of the two cases above. Isotopic Fractionation Induced by Photosynthesis: For more info: John Hayes chapter in Valley and Cole (Ch. 3). Fogel and Cifuentes (1993) Isotope fractionation during primary production, in Organic Geochemistry: Principles and Applications. M.H. Engel and S.A. Macko, eds., pp. 7398, Plenum, New York.

3 There are 3 distinct photosynthesis metabolisms (C 3, C 4, and CAM) with contrasting C isotope fractionation: 1) C 3 Plants: Calvin cycle: Algae, autotrophic bacteria, most land plants For more info See Fogel and Cifuentes, 1993 Fig.1 for sketch of RuBP reaction and their Fig. 3 for C isotope fractionation by plants as a function of CO 2 conc. Picture an algal cell, with CO 2 diffusing in, and some CO 2 inside diffusing out. CO 2 (out) CO 2 (in) Organic C RuBP enzyme CO 2 (outside) Step 1: Diffusion; Kinetic, ε < 4.4 CO 2 (inside) Step 2: Carboxylation via RuBP (Enzyme); Kinetic, ε = 29.4 Phosphoglyceric acid (3 Carbons) Later steps (no kinetic isotope effect) What is the resulting overall fractionation? Case 1: If step 2 is ratelimiting: ε total = ε 1 + ε 2 30 When does this occur? 1) When CO 2 is so plentiful that CO 2 concentration inside the cell is affected little by photosynthetic uptake; diffusion is not limiting. 2) OR When photosynthesis is slow, so diffusion is, in comparison, very fast. Case 2: If step 1 is ratelimiting: ε total = ε 1 (small) When does this occur? 1) When CO 2 is not plentiful, and the limited CO 2 inside the cell is scavenged strongly by photosynthesis so the concentration there is drawn down. 2) OR When photosynthesis rate is very high and becomes limited by resupply of CO 2 by diffusion. Case 3: In most cases, diffusion DOES limit the rate SOMEWHAT. This case is intermediate between cases 1 and 2 and the fractionation is also intermediate.

4 The extent to which diffusion is ratelimiting depends on the CO 2 concentration inside, relative to outside. So one way to model ε is as a function of C in /C out. Here is model for the process: ε = a + (c i /c a )(ba) a = step 1 fract. b = step 2 fract. c I = CO 2 conc. inside c a = CO 2 conc. outside (White uses Δ instead of ε) 30 ε ( ) Theoretical line: ε = (C in /C out ) (seems like real plants have somewhat smaller fractionations than predicted by theory, but the general trend is correct and values are similar) 0 0 C in /C out Data from real C 3 plants 1.0 Generally observed ε values for C 3 plants: 16 to 26 Depends on all variables affecting plant s rate of photosynthesis Faster photosynthesis rates mean lower conc. inside Generally observed δ 13 C for C 3 plants: 23 to 33 (atmospheric CO 2 δ 13 C = 7 ) IMPORTANT APPLICATION: We expect ε decreases when CO 2 conc. in the atmosphere decreases. Thus, if we can determine ε for photosynthesis at some time in the past and compare it to today s value we may have an indication of past CO 2 concentrations. Also: with higher plants, we see that water stress causes plants to close stomatae and then to have less strong fracttionation

5 2) C 4 Plants:: Sugar cane, Maize/corn, hotregion grasses; these plants were later to evolve than C 3 plants, and proliferated in midlatitude areas relatively recently. Whereas the C 3 plants do not manage CO 2 flows into and out of the cells much, C 4 plants have a special mechanism that grabs CO 2 and shuttles it to a place where it cannot escape so easily. Diffusion ε = 0.8 CO 2 (out) CO 2 (in) Step 1: Diffusion; Kinetic, ε = 4.4 Step 2: CO 2 /HCO 3 equilibrium, Δ = +8 HCO 3 Oxaloacetate (4 C s) Shuttle molecule Oxaloacetate (4 C s) inside sheath cells Step 3: Fixation via PEP (Enzyme); Kinetic, ε = 2 Step 4: Transport to inner bundle sheath cells and conversion back to CO 2 ; Kinetic, ε = 0? CO 2 Leakage CO 2 Step 5: Carboxylation via RuBP (Enzyme); Kinetic, ε = 29.4 Phosphoglyceric acid (3 C s) Later steps (no effect) A model equation for C 4 plants is (version in White s notes) ε = a +(b b 5 φ a) (c i /c a ) (White uses Δ instead of ε) a = step 1 fract. b 234 = steps (2+3+4) fract. b 5 = step 5 fract. (this is the large ε step) φ = fraction of CO 2 leaked to outside from the inner sheath cells A more complete equation for C 4 plants is (according to Hayes; better reference) ε = a +[b 3 b 2 + φ(b 5 b leak ) a ](1f) a = step 1 fract. (diffusion, atm into plant; ε = 4.4 ) b 2 = step 2 fract. (equilibrium between CO 2 and bicarbonate, both inside; Δ =+8 ) b 3 = step 3 fract. (PEP binding of bicarb, ε = 2.2 ) b 5 = step 5 fract. (RuBP reaction, ε = 25 ) b leak = fract. occurring with leakage of CO 2 out of sheath cells φ = fraction of CO 2 leaked to outside from the inner sheath cells (varies between plants, but f is the fraction of carbon taken into the plant that actually makes it through the reaction

6 Note that (1f) can be related to CO 2 concentrations inside the plant and outside the plant as before (inside meaning inside the plant but not inside the sheath cells yet) ε = a +[b 3 b 2 + φ(b 5 b leak ) a ](c i /c a ) The bottom line: Average ε for C 4 plants is about 4 Mean δ 13 C for C 4 plants is close to 11 (atm CO 2 = 7 ) (versus C 3 plants = 20 to 30 ) C 4 plants are also less sensitive to changes in CO 2 conc. in the atmosphere. The general picture: C 4 plants shuttle CO 2 into inner cells that are separated from the outside world. The carboxylation step that causes the strongest fractionation is not much in contact with the outside environment, and thus the overall fractionation for photosynthesis is small. 30 Data from real C 3 plants 0 φ =0.34 line Data from real C 4 plants φ =0.21 line 3) CAM plants (Mostly cacti and other succulents) Similar to C 4 plants, in that they manage the flow of CO 2 and its consumption by photosynthesis. Night: Stomata open, CO 2 is taken in and held by a carrier molecule Day: Stomata close. CO 2 released to react with RuBP (Highly advantageous for desert plants to keep their stomata closed during the day saves water) Result: ε = 4 to 26 (low end is for very tight plants; little leakage during day) 4) Marine algae: a. Use HCO 3 instead of CO 2, usually Less fractionation (relative to CO 2 supply) than land plants, because the CO 2 HCO 3 equilibrium involves an equilibrium fractionation the HCO 3 is heavier by 7 to 12 b. Diffusion is slower in water than in air, so greater tendency to be diffusionlimited so fractionation strongly dependent on the CO 2 (actually, DIC = dissolved inorganic carbon = total of CO 2 + HCO 3 + CO 3 = ) concentration. Marine algae have pump mechanisms that actively transport DIC into cell, thus cell conc. is higher than outside. ε = d +b 3 (F 3 /F 1 ) (Note: White has a typo in his version) d = step 1 fract. equilibr. Between CO 2 + HCO 3 (Δ = +8.5 ) b 3 = fract for carboxylation (29 ) F 3 /F 1 = fraction of the CO 2 leaked out of the cell Summary: ε for marine algae is variable; somewhat smaller than for land plants; most common range: 1015

7 5) Fresh water algae: Somewhat like marine algae, but often ph is lower (vs. ocean = 7.5 to 8.5), HCO 3 utilization less important; fractionation somewhat greater as a results Equilibrium fractionations for aqueous DIC species Species Δ vs. CO 2 (gas) at 25 C CO 2 (dissolved) 0.8 HCO = CO CaCO These are, of course temperature dependent. These values are just to give you a rough idea of direction and magnitude of fractionation. Biosynthesis additional, smaller C isotope fractionation between different molecules that are manufactured from the products of photosynthesis (and fractionation between differently bonded carbons in a single molecule) This process fractionates isotopes somewhat (<10 ), so different components of an organism will have different δ 13 C values. lipids tend to be lighter than other components by 2 to 8 general tendency to have lighter C in CH bonded groups (lipids! petroleum) general tendency to have heavier C in C0 bonded groups (sugars, cellulose, amino acids) makes some sense in terms of equilibrium theory (but these reactions are kinetic). Here is a figure from Hoefs (1987) giving ranges of δ 13 C for autotrophs Carbon isotopes on Earth 1. Ranges of various carbon reservoirs on earth: (δ 13 C vs. PDB (or VPDB)) Mantle: 3 to 7 Atmosphere ~ 7 Terrestrial Plants ~ 8 to 35 Marine organic matter in seds. ~ 15 to 35 Seawater HCO 3 = ~ 2 Marine Carbonate precipitated ~ 0 Methane ~ 10 to 80 NOTE: Atmospheric CO 2 δ 13 C has changed over time. For example, fossil fuel combustion has released C with a relatively low δ 13 C value into the atmosphere, causing a shift in δ 13 C (about 0.2 between 1982 and 1990, for example). Any disturbance in the global C cycles will tend to change the atmospheric δ 13 C value.

8 C isotope fractionation during production of methane: 3 pathways for creation of methane: 1. Biogenic gas; Fermentation: Methane is generated by microbes breaking organic molecules into smaller ones; e.g., acetate cleaved into CH 4 + CO 2 C source has low δ 13 C AND methane production adds a kinetic isotope effect, so Methane produced is very light isotopically 2. CO 2 reduction: CO 2 is a terminal electron acceptor for bacteria methane produced, and it is much lighter than the starting CO 2, resulting in highly negative δ 13 C values. Usually in marine sediments. 3. Thermogenic gas: Abiotic breakage of molecules; different fractionation than fermentation

9 These three types of gas can be distinguished by C and H isotopes: 100 δ 13 C Fermentation CO 2 reduction Thermal δd 100 See Whiticar et al., 1986, Geochim. Cosmichim. Acta, 50, Some applications of C isotopes (there are MANY others): Because carbon is central to biological processes AND it undergoes large kinetic and/or equilibrium isotope shifts during those processes, variations in C isotopes in many different materials often hold information about biological processes and the global and local carbon cycles, and how they have changed in earth history. Marine sediments and sedimentary rocks provide excellent archives but only if we can interpret what the various δ 13 C records mean. Note that because the carbon cycle has various inputs and outputs, the δ 13 C of the various materials and pools of carbon can only be understood if we consider the response of each system to changes in its inputs and outputs. This can be a difficult task, but it is made easier if we can approximate the actual system by one that is in steady state, meaning the system inputs and outputs are balanced. Application #1: Impact of fossil fuels on CO2 budget of the atmosphere. δ 13 C in atmospheric CO 2 has decreased since the beginning of the industrial revolution (see figure in White s notes). This is consistent with the idea that the atmosphere has received large amounts of isotopically light CO 2 from the burning of fossil fuels. However, if you look in detail you see that the shift is not all that large, and reflect greater complexity (probably rapid cycling of C through biomass and the surface layer of the ocean). Application #2: Paleoproductivity of the oceans. There MUST be some positive feedback that causes the earth system to have an amplified response to the rather weak Milankovitch forcing. One possible feedback mechanism is nutrient supply changes that affect the strength of the biological pump that removes CO 2 from the atmosphere and makes marine phytoplankton (algae). Greater phytoplankton productivity should suck more isotopically light DIC from the water, and the remaining DIC in the surface layer should become isotopically heavier. Measure: δ 13 C in planktonic (surfacedwelling) forams Problem: The entire ocean may change over time, so we need to correct for that. Solution: Measure benthic forams to get δ 13 C for deep ocean Result: Differences in δ 13 C do indeed suggest changes in productivity that correlate with glacial cycles. Application #3: Determining past pco 2 in the atmosphere, going back in time beyond the extent of ice core records. The magnitude of C isotope fractionation during photosynthesis by marine algae depends on CO 2 concentration in the water (which is linked to that in the atmosphere). We see changes in δ 13 C as the CO 2 conc. changes from glacial to interglacial.

10 " One approach is to measure ALL the organic molecules extracted from sediments. This seems to work: See Figures in White notes: " That approach can suffer from changes in the supply of organic matter and from its degradation. It is better to pick certain molecules that are derived from marine algae only and are resistant to degradation. " References for these approaches: Kate Freeman s chapter in the Valley and Cole Book describes this. Jasper J. P. and Hayes J. M. (1994) Reconstruction of paleoceanic pco 2 levels from carbon isotopic compositions of sedimentary biogenic components. In Carbon Cycling in the Glacial Ocean: Constraints on the Ocean's Role in Global Change (ed. T. F. P. R. Zahn, M.A. Kaminski and L. Labeyrie), pp SpringerVerlag. Rau G. (1994) Variations in sedimentary organic δ 13 C as a proxy for past changes in ocean and atmospheric CO2 reconstructions. In Carbon Cycling in the Glacial Ocean: Constraints on the Ocean's Role in Global Change (ed. T. F. P. R. Zahn, M.A. Kaminski and L. Labeyrie), pp Springer Verlag. Example: Jasper and Hayes (1994) used the following relationship between the C isotope fractionation during manufacture of certain organic molecules and the dissolved CO 2 concentration: ε = 32.9 log(c) (C is micromolar dissolved CO 2 in ocean) Example 2: Late Precambrian glaciations, around 750 Ma: Negative excursions in δ 13 C, along with decreases in ε. Smaller fractionation implies low CO 2 in atmosphere, and resulting cold climate have been the driver of the glaciations. Application #3: Over the entire 4 billion year history of the earth, major changes in the earth s carbon cycle are reflected in carbon isotopes in ocean seds. See DesMarais chapter in Valley and Cole if you want more info. C inputs to Oceans; δ 13 C 5 Oceans OrgC burial Carbonate burial ε 20 ε 5 Use a steadystate model (reasonable approximation); inputs are balanced by outputs Organic C vs. carbonate burial i. Organic C burial removes light C, drives ocean DIC to greater δ 13 C ii. Carbonate burial removes heavy C, drives ocean DIC to lesser δ 13 C iii. If organic C burial dominates, ocean DIC quite heavy ( > +20 ) iv. If carbonate burial dominates, ocean not so heavy δ 13 C carbonate organic C Fig. 3 of Des Marais chapter from Valley and Cole Book. Solid lines give composition of the 2 major outputs of C from the oceans, as functions of the % of total burial occurring as organic C burial % of C burial as organic C 100% Currently, 20% of the carbon burial is organic C

11 Before oxygenic photosynthesis: i. Organic carbon synthesis small; only about 20 Teramoles per year ii. CO 2 outgassing from mantle about the same size After oxygenic photosynthesis: i. Organic carbon synthesis about 9000 Teramoles/yr. ii. What happens if this is mostly buried? Changes on earth at about 2.2 Ga: i. Huge swings to greater δ 13 C in the ocean s DIC ii. Presumably reflects greater burial of organic matter iii. Oxygenic photosynthesis cranked up production of organics iv. Mostly buried until scavenging organisms evolved to be more efficient Negative excursions in δ 13 C of carbonate around 750 Ma lower productivity during snowball earth conditions? Little organic burial. Application #4: Similarly: C isotope profiles across the KT boundary carbonate rocks record a decrease in δ 13 C just after the KT boundary Probably reflects strong changes in global C cycle after a meteorite impact We expect lower productivity, therefore, less burial of isotopically light organic C! This burial normally drives DIC toward heavier values! Less of this means lighter values Application #5: The diet of animals and hominids can be determined from the δ 13 C (and δ 15 N) values. Food remains in pot shards are one possibility for something to measure, teeth and bones are possible for δ 13 C (see White s Chapter on this topic) Legumes have distinctive N isotope ratios Maize is a C 4 plant shifts observed in anthropological sites Diets of marine species can also be traced (Paul Koch, UC Santa Cruz and others) Birds, too. Application #6: Also, δ 13 C is influenced by variations in biological productivity in water. δ 13 C in DIC tends to be a function of depth in a body of water where photosynthesis occurs in the top layer. Lighter isotopes removed preferentially, converted to solid particles. If productivity was less in the past (maybe because of climate change or nutrient flux change) we expect that the δ 13 C will reflect this. (Productivity of the oceans is critical in the global carbon cycle if productivity declines, less C is sucked out of the atmosphere.) Application #7: Degradation of organic groundwater contaminants often involves a kinetic isotope effect. The remaining contaminant becomes progressively heavier with increasing degradation. This provides a means of determining degradation rates. Example: MTBE degradation: ε = 15.6 ± 4.1 in laboratory (Somsamak et al., ES&T vol. 39, , 2005) If there are multiple sources of pollution in one area, it may be possible to fingerprint the source of contaminants with δ 13 C measurements (and δd). δ 37 Cl may also provide an indicative tracer for Clbearing compounds like TCE. Measurement of C isotopes δ 13 C is measured by gas source mass spectrometers via 3 different approaches: Method #1: Offline preparation. Generate CO 2 from your sample, measure by simple injection of the gas into the mass spec source, alternated with standard CO 2

12 Method #2: Compoundspecific δ 13 C. Separate various organic molecules using a gas chromatograph. Pass the compound you want to analyze into a combustion furnace, where it is converted to CO 2 or CO. The small pulse of gas is passed into the ionization chamber of the mass spec and gives a brief, transient signal (harder to measure well, precision not as good at Method #1). This method can be very helpful because sediments usually contain many molecules of various origin and type. Isolating one molecule that originated from a single process is often necessary to avoid mixed signals and concentrate on a single shift in nature. Method #3: Elemental Analyzer. Combust a sample at high T in a special furnace that generates CO. Send the transient pulse of gas to the mass spec. Basically, any method that generates a carbonbearing gas and shifts the C isotopes consistently (or not at all) can be used. δ 13 C can be measured by infrared absorption spectrometry also. Small, portable, inexpensive instruments; good for field measurements, but may be less precise.

Carbon Isotopes and Isotopic Fractionation for Complex Processes

Carbon Isotopes and Isotopic Fractionation for Complex Processes Reading: Valley and Cole Chapters 3 (Hayes),10 (Des Marais), and 11 (Freeman) White, Lecture 28 Optional: Faure, 1986, Chapter 26 Motivation: