Stable Isotopes
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STABLE ISOTOPES Dr.Thrivikramji.K.P [email protected] Introduction Isotopes are atoms of same element having different numbers of neutrons leading to differing atomic masses. With the exception of C, general geological processes do not affect separation of the isotopes used in absolute dating. But certain elements of
low atomic no. tend to show very systematic variations in isotope ratios in very ordinary geological processes like weathering, crystallization, diagenesis, authigenesis etc. By precise measurements of isotope ratios, inferences can be drawn about the conditions and reactions that created such changes. The variabilities are a result of different frequencies of vibration of heavy and light atoms in a molecule or crystal structure. Compared to heavy isotopes, atoms of light isotopes vibrate with higher frequencies and so are bonded to others less strongly. Process- dependent separation of heavy and light isotopes is characteristic of elements of low atomic no. Differing vibrational frequencies of particles diminish with increasing temp. resulting in much less pronounced separation. Table 1. Important Elements and isotopes Element Oxygen Hydrogen Carbon Sulfur
Isotopes 16 O, 17O and 18O (17O least abundant) 1 H & 2H (D), 3H (T radio ctive) 12 C, 13C and 14C (radio active) 32 33 S, S, 34S and 36S (only 32 &34 studied)
Chief mechanisms of isotope fractionation in nature are: a. Physical properties controlled mechanisms like evaporation and diffusion. During evaporation of water light isotopes of O and H enter the vapor phase leaving the heavier isotopes in the liquid water. b. In equilibrium reactions of the following type: ½ C16O2 +H2 18O
½
18
CO2 + H216O
c. Separations based on reaction rates. In bacterial reduction of sulfate production of sulfide is faster for the lighter isotope (32S) than for the heavy one with the result 34S is enriched in the residual phase. Extent of separation is defined by fractionation factor, = Ra/Rb where, Ra and Rb are ratios of concentration of heavy to light isotope in phases a and b and at equilibrium, for water and vapour, is 1.0092
1
As such numbers are small and hence hard to visualize, another scheme of symbolism is used to describe isotope separation i.e., heavy = [(heavy/light in sample) (heavy/light in standard)] / [(heavy/light in standard )] x 1000 per mil and for oxygen it can be re-written as: 18
O = [( 18O/16O) spl (18O/16O) std] / [(18O/16O) std)] x 1000
heavy or 18O can assume a +ve or ve value meaning the sample is rich in heavy isotope or enriched in light isotope respectively. The isotopic compositions of materials analyzed on mass spectrometers are usually reported relative to some international reference standard. Samples are either analyzed at the same time as this reference standard or with some internal laboratory standard that has been calibrated relative to the international standard. Alternatively, the absolute ratios of isotopes can be reported. Small quantities of these reference standards are available for calibration purposes from either the National Institute of Standards and Technology (NIST) in the USA . Various isotope standards are used for reporting light stable-isotopic compositions. The d values of each of t he st andards have been defined as 0 . D and 18O values are normally reported relative to the SMOW standard (Standard Mean Ocean Water; Craig, 1961) or the equivalent VSMOW (Vienna-SMOW) standard. 13C values are reported relative to either the PDB (Pee Dee Belemnite) or the equivalent VPDB (Vienna-PDB) standard. 18O values of low-temperature carbonates are also commonly reported relative to PDB or VPDB.
Fractionation of isotopes of elements O and H is a very important phenomenon in nature. In comparison to standard, light isotopes of O and H are enriched in the fresh water, glacial ice and snow on the continents and have 18O and D values. Process of enrichment of light isotope increases with decrease in air temp and hence varies seasonally and latitudinally. Salinity showed a positive correlation with isotope values of sea water and hence offered an insight into the question of origin of CBW in oceans. Past oceanic temperatures have been estimated by the isotope composition of carbonate, silica and phosphate. It suggests that climatic temp had declined from 70o C at 3.4 Ga to present values and were between 34-20 oC in Paleozoic era. Continental ice sheet in Antarctica in Miocene epoch showed an increase in 18 O in benthic foraminifera. In these estimates it is assumed that isotopic equilibrium existed between the sea water and mineral phase in the organism. Isotope composition of land snails (carbonate), speleothems (carbonate) and mammalian and fish bones (phosphate) have been used to estimate past
2
continental temps. More over isotopic composition of meteoric water came in handy to examine the nature and origin of geothermal water and brine. As O being a major constituent of rocks, when two minerals A&B equilibrate with isotopes in a common reservoir, the difference in the 18O values decrease with increasing temp. This is the basis of determination of final equilibrium temp of cogenetic minerals having O. Generally minerals of volcanic origin are better equilibrated at the time of formation in the melt and hence give better temp of crystallization than plutonic rocks where due to re-equilibration of O as a result of slow cooling. Further in plutons recirculating meteoric waters also by lowering the 18 O of minerals while raising that of water. Re-equilibaration of O isotope is accompanied by deuteric alteration of feldspar and other minerals. Waters of different origins may be involved in the formation of ores and wall rock alteration and hence a complex picture of isotope composition is the result. S type granites are enriched in 18 O than the I-types granites. Temp and type of water determine the isotopic composition of O and H in clay minerals. So equations can be devised to show the D and 18O relationship with clay species and meteoric water, and hence the climate at he tie of their formation. Isotopic composition can be also used to discriminate the supergene (high temp) and hypogene (low temp) cays in ore deposits. Marine carbonates have +ve 18O values from +20 permil to +30 permil relative to SMOW. 18O values of cherts and carbonates decrease with increasing age of 3.5Ga. This nature is variously explained as below: 1. isotopic re-equilibration with meteoric water, 2) a decrease in temp of ocean water from archaean to present. 3, an increase in 18O values since archaean. It is noticed that differences in 18O values of minerals metamorphic rocks generally decrease with increasing grade of metamorphism. Permeability of rocks determine the extent of re-equilibration by exchange with aqueous fluid or CO2 during metamorphism. In the absence of permeability large isotopic differences will be noticed over short distances. Carbon (Z=6) Basis of all life on earth and one of the most abundant element in the universe Fractionation by inorganic exchange reactions and by photosynthesis of green plants. Carbonates enriched in 13C while plants and animals are enriched in 12C indicating the derivation from biogenic matter. CO2 levels in the atmosphere have gone up by 10% and along with 12C due to burning of fossil fuels in the last century. Reduced C in sedimentary rocks of PC age are rich in 13C indicating a biogenic origin.
3
Isotopes are fractionated in the system CO2 (gas) - carbonate ions aqueous CaCO3 (solid) so that calcite is enriched in 13C by 10 permil at 20 deg C relative to CO2 gas. 13C values are closer to zero relative to PDB and also do not vary appreciably by age. Carbonates of lacustrine derivation are slightly enriched in 12C due to derivation of C from decaying plant debri of soil. Carbonaceous chondrites are enriched in 13C while carbonates of terrestrial origin are rich in 12C formed by oxidation of bacteriogenic CH4. There is a slight enrichment 13C between the Precambrian and phanrozoic rocks. 13C of carbonates and diamonds vary significantly either due to isotopic heterogeneity of C in the mantle or due to fractionation or both. Graphite is variably enriched in 12C compatible with its biogenic source, whereas 13C rich graphite may be abiogenic. Isotopic profile of carbonate minerals CO2 gas in fluid inclusions are variable suggesting deep seated sources. Sulfur (Z= 16) Dist ribut ed in all t he eart h s spheres including biosphere and occurs in reduced (in metallic deposits and sulfide ore bodies associated with igneous, sedimentary and metamorphic rocks) , native (salt domes) and oxidized states (in ocean and evaporites). S has four stable isotopes, viz., 32S (=95.02%), 33S (=0.75), 34S (=4.21%) and 36S (= 0.02%) and isotopic composition is defined as 34 S. Fractionation of S, during bacterial (Desulfovibrio desulfuricanus) reduction of sulphate to sulfide and by isotope exchange reactions among sulfur bearing compounds, ions and molecules are mechanisms enrich hydrogen sulfide with 32 S by 50% or more. In lab cultures only 27% enrichment was observed. Extent of enrichment depends on factors like rate of reduction, temperature, nature and availability of food supply and size of sulfate reservoir. Due to role of bacteria, in modern environments H2S and sulfide minerals are variously enriched in 32S. 34S values of petroleum range between -8 to +32 permil and appear to be enriched in 32S by about 15% in comparison with marine sulfate at the time of formation petroleum. 34S values in sulfur of coal are nearly constant which is considered to be due to variability of marine sulfate and its constancy in fresh water. Isotopic composition of C and S in cap rock of salt domes is explained as follows. Native S in saltdomes formed as a byproduct of sulfate reduction by bacteria feeding on petroleum. Cap rock calcite is enriched in 12C due to precipitation from metabolic CO2, again derived from petroleum.
4
Bacterial sulfate reduction became important only from 2.5Ga, after photosynthetic sulfur reducing bacteria emerged and increased sulfate reduction oceanic waters. Grenville (Ontario) evaporates strongly suggest activity of sulfate reduction bacteria in ocean during 1000-1300 Ma an in similar rocks of Zimbabwe. Isotopic compositions of sulfate minerals igneous rocks are similar to S of Canyon Diablo meteorite. Granitic rocks showed variable S values either due to magma forming by melting of sediments or magma got contaminated by biogenic S of crust. Use of S isotopes of sulfide ore bodies to model the origin of igneous hydrothermal from sedimentary syngenetic ones failed due to extensive overlap of 34S values. Sulfur entering the atmosphere from natural (volcanic sources) and anthropogenic (combustion and refining of fossil fuels) sources, have a range of compositions, and truly create an environmental hazard. -----------------------Suggested reading Calvin, M., and A. A. Benson. 1948. The path of carbon in photosynthesis. Science v.107 pp.476-80. Coleman, David, and Brian Fry. 1991. Carbon Isotope Techniques. Academic Press/Harcourt Brace Jovanovich, New York. De Niro, M. J., and S. Epstein. 1978. Influence of diet on the distribution of carbon isotopes in animals. Geochim. et Cosmochim. Acta v.42 pp.495-506. De Niro, Michael J. 1987. Stable Isotopy and Archaeology. American Scientist v.75 pp.182-187. Ehleringer, J. R. and P.W. Rundel. 1989. Stable Isotopes: History, Units, and Instrumentation. In Rundel, P. W., J. R. Ehleringer and K. A. Nagy, eds. Stable Isotopes in Ecological Research. Springer Verlag, New York. Hall, R. L. 1967. Those late corn dates: Isotopic fractionation as a source of error in carbon-14 dates. Mich. Archaeol. v.13 pp.171-79. Hatch, M. D. and C. R. Slack. 1970. The C4 carboxylic acid pathway of photosynthesis. pp.35-106. In Reinhold L and Liwschitz Y, eds., Progress in Phytochemistry. Wileylnterscience, New York. Hatch, M. D., and C. R. Slack. 1966. Photosynthesis by sugarcane leaves. A new carboxylation reaction and the pathway of sugar formation. Biochem. J. v.101 pp.10311. Hayes, J. M. 1982. Fractionation et. al.: An introduction to isotopic measurements and terminology. Spectra v.8 pp.3-8. Rounick, J. S. and M. J. Winterbourn. 1986. Stable carbon isotopes and carbon flow in ecosystems. BioScience v.36 pp.171-176. Rundel, P. W., J. R. Ehleringer and K. A. Nagy (eds.) 1989. Stable Isotopes in Ecological Research. Springer Verlag, New York. Smith, B. N. and S. Epstein. 1971. Two categories of 13C/12C ratios for higher plants. Plant Physiol. v.47 pp.380-384.
5
van der Merwe, Nikolaas J. 1982. Carbon isotopes, Photosynthesis, and Archaeology. American Scientist, Nov-Dec 1982, 596-606. van der Merwe, N. J., and J. C. Vogel. 1978. 13C content of human collagen as a measure of prehistoric diet in Woodland North America. Nature v.276 pp.815-16. van der Merwe, Nikolaas J. 1982. Carbon isotopes, photosynthesis, and archaeology. American Scientist v.70 pp.596-605. Vogel, J. C. 1978b. Isotopic assessment of the dietary habits of ungulates. South Afr. J. Sci. v.74 pp.298-301. Vogel, J. C. and N. J. van der Merwe. 1977. Isotopic evidence for early maize cultivation in New York State. Am. Antiquity v.42 pp.238-42. Zelitch, I. 1971. Photosynthesis, Photorespiration, and Plant Productivity. Academic Press, New York.
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6
low atomic no. tend to show very systematic variations in isotope ratios in very ordinary geological processes like weathering, crystallization, diagenesis, authigenesis etc. By precise measurements of isotope ratios, inferences can be drawn about the conditions and reactions that created such changes. The variabilities are a result of different frequencies of vibration of heavy and light atoms in a molecule or crystal structure. Compared to heavy isotopes, atoms of light isotopes vibrate with higher frequencies and so are bonded to others less strongly. Process- dependent separation of heavy and light isotopes is characteristic of elements of low atomic no. Differing vibrational frequencies of particles diminish with increasing temp. resulting in much less pronounced separation. Table 1. Important Elements and isotopes Element Oxygen Hydrogen Carbon Sulfur
Isotopes 16 O, 17O and 18O (17O least abundant) 1 H & 2H (D), 3H (T radio ctive) 12 C, 13C and 14C (radio active) 32 33 S, S, 34S and 36S (only 32 &34 studied)
Chief mechanisms of isotope fractionation in nature are: a. Physical properties controlled mechanisms like evaporation and diffusion. During evaporation of water light isotopes of O and H enter the vapor phase leaving the heavier isotopes in the liquid water. b. In equilibrium reactions of the following type: ½ C16O2 +H2 18O
½
18
CO2 + H216O
c. Separations based on reaction rates. In bacterial reduction of sulfate production of sulfide is faster for the lighter isotope (32S) than for the heavy one with the result 34S is enriched in the residual phase. Extent of separation is defined by fractionation factor, = Ra/Rb where, Ra and Rb are ratios of concentration of heavy to light isotope in phases a and b and at equilibrium, for water and vapour, is 1.0092
1
As such numbers are small and hence hard to visualize, another scheme of symbolism is used to describe isotope separation i.e., heavy = [(heavy/light in sample) (heavy/light in standard)] / [(heavy/light in standard )] x 1000 per mil and for oxygen it can be re-written as: 18
O = [( 18O/16O) spl (18O/16O) std] / [(18O/16O) std)] x 1000
heavy or 18O can assume a +ve or ve value meaning the sample is rich in heavy isotope or enriched in light isotope respectively. The isotopic compositions of materials analyzed on mass spectrometers are usually reported relative to some international reference standard. Samples are either analyzed at the same time as this reference standard or with some internal laboratory standard that has been calibrated relative to the international standard. Alternatively, the absolute ratios of isotopes can be reported. Small quantities of these reference standards are available for calibration purposes from either the National Institute of Standards and Technology (NIST) in the USA . Various isotope standards are used for reporting light stable-isotopic compositions. The d values of each of t he st andards have been defined as 0 . D and 18O values are normally reported relative to the SMOW standard (Standard Mean Ocean Water; Craig, 1961) or the equivalent VSMOW (Vienna-SMOW) standard. 13C values are reported relative to either the PDB (Pee Dee Belemnite) or the equivalent VPDB (Vienna-PDB) standard. 18O values of low-temperature carbonates are also commonly reported relative to PDB or VPDB.
Fractionation of isotopes of elements O and H is a very important phenomenon in nature. In comparison to standard, light isotopes of O and H are enriched in the fresh water, glacial ice and snow on the continents and have 18O and D values. Process of enrichment of light isotope increases with decrease in air temp and hence varies seasonally and latitudinally. Salinity showed a positive correlation with isotope values of sea water and hence offered an insight into the question of origin of CBW in oceans. Past oceanic temperatures have been estimated by the isotope composition of carbonate, silica and phosphate. It suggests that climatic temp had declined from 70o C at 3.4 Ga to present values and were between 34-20 oC in Paleozoic era. Continental ice sheet in Antarctica in Miocene epoch showed an increase in 18 O in benthic foraminifera. In these estimates it is assumed that isotopic equilibrium existed between the sea water and mineral phase in the organism. Isotope composition of land snails (carbonate), speleothems (carbonate) and mammalian and fish bones (phosphate) have been used to estimate past
2
continental temps. More over isotopic composition of meteoric water came in handy to examine the nature and origin of geothermal water and brine. As O being a major constituent of rocks, when two minerals A&B equilibrate with isotopes in a common reservoir, the difference in the 18O values decrease with increasing temp. This is the basis of determination of final equilibrium temp of cogenetic minerals having O. Generally minerals of volcanic origin are better equilibrated at the time of formation in the melt and hence give better temp of crystallization than plutonic rocks where due to re-equilibration of O as a result of slow cooling. Further in plutons recirculating meteoric waters also by lowering the 18 O of minerals while raising that of water. Re-equilibaration of O isotope is accompanied by deuteric alteration of feldspar and other minerals. Waters of different origins may be involved in the formation of ores and wall rock alteration and hence a complex picture of isotope composition is the result. S type granites are enriched in 18 O than the I-types granites. Temp and type of water determine the isotopic composition of O and H in clay minerals. So equations can be devised to show the D and 18O relationship with clay species and meteoric water, and hence the climate at he tie of their formation. Isotopic composition can be also used to discriminate the supergene (high temp) and hypogene (low temp) cays in ore deposits. Marine carbonates have +ve 18O values from +20 permil to +30 permil relative to SMOW. 18O values of cherts and carbonates decrease with increasing age of 3.5Ga. This nature is variously explained as below: 1. isotopic re-equilibration with meteoric water, 2) a decrease in temp of ocean water from archaean to present. 3, an increase in 18O values since archaean. It is noticed that differences in 18O values of minerals metamorphic rocks generally decrease with increasing grade of metamorphism. Permeability of rocks determine the extent of re-equilibration by exchange with aqueous fluid or CO2 during metamorphism. In the absence of permeability large isotopic differences will be noticed over short distances. Carbon (Z=6) Basis of all life on earth and one of the most abundant element in the universe Fractionation by inorganic exchange reactions and by photosynthesis of green plants. Carbonates enriched in 13C while plants and animals are enriched in 12C indicating the derivation from biogenic matter. CO2 levels in the atmosphere have gone up by 10% and along with 12C due to burning of fossil fuels in the last century. Reduced C in sedimentary rocks of PC age are rich in 13C indicating a biogenic origin.
3
Isotopes are fractionated in the system CO2 (gas) - carbonate ions aqueous CaCO3 (solid) so that calcite is enriched in 13C by 10 permil at 20 deg C relative to CO2 gas. 13C values are closer to zero relative to PDB and also do not vary appreciably by age. Carbonates of lacustrine derivation are slightly enriched in 12C due to derivation of C from decaying plant debri of soil. Carbonaceous chondrites are enriched in 13C while carbonates of terrestrial origin are rich in 12C formed by oxidation of bacteriogenic CH4. There is a slight enrichment 13C between the Precambrian and phanrozoic rocks. 13C of carbonates and diamonds vary significantly either due to isotopic heterogeneity of C in the mantle or due to fractionation or both. Graphite is variably enriched in 12C compatible with its biogenic source, whereas 13C rich graphite may be abiogenic. Isotopic profile of carbonate minerals CO2 gas in fluid inclusions are variable suggesting deep seated sources. Sulfur (Z= 16) Dist ribut ed in all t he eart h s spheres including biosphere and occurs in reduced (in metallic deposits and sulfide ore bodies associated with igneous, sedimentary and metamorphic rocks) , native (salt domes) and oxidized states (in ocean and evaporites). S has four stable isotopes, viz., 32S (=95.02%), 33S (=0.75), 34S (=4.21%) and 36S (= 0.02%) and isotopic composition is defined as 34 S. Fractionation of S, during bacterial (Desulfovibrio desulfuricanus) reduction of sulphate to sulfide and by isotope exchange reactions among sulfur bearing compounds, ions and molecules are mechanisms enrich hydrogen sulfide with 32 S by 50% or more. In lab cultures only 27% enrichment was observed. Extent of enrichment depends on factors like rate of reduction, temperature, nature and availability of food supply and size of sulfate reservoir. Due to role of bacteria, in modern environments H2S and sulfide minerals are variously enriched in 32S. 34S values of petroleum range between -8 to +32 permil and appear to be enriched in 32S by about 15% in comparison with marine sulfate at the time of formation petroleum. 34S values in sulfur of coal are nearly constant which is considered to be due to variability of marine sulfate and its constancy in fresh water. Isotopic composition of C and S in cap rock of salt domes is explained as follows. Native S in saltdomes formed as a byproduct of sulfate reduction by bacteria feeding on petroleum. Cap rock calcite is enriched in 12C due to precipitation from metabolic CO2, again derived from petroleum.
4
Bacterial sulfate reduction became important only from 2.5Ga, after photosynthetic sulfur reducing bacteria emerged and increased sulfate reduction oceanic waters. Grenville (Ontario) evaporates strongly suggest activity of sulfate reduction bacteria in ocean during 1000-1300 Ma an in similar rocks of Zimbabwe. Isotopic compositions of sulfate minerals igneous rocks are similar to S of Canyon Diablo meteorite. Granitic rocks showed variable S values either due to magma forming by melting of sediments or magma got contaminated by biogenic S of crust. Use of S isotopes of sulfide ore bodies to model the origin of igneous hydrothermal from sedimentary syngenetic ones failed due to extensive overlap of 34S values. Sulfur entering the atmosphere from natural (volcanic sources) and anthropogenic (combustion and refining of fossil fuels) sources, have a range of compositions, and truly create an environmental hazard. -----------------------Suggested reading Calvin, M., and A. A. Benson. 1948. The path of carbon in photosynthesis. Science v.107 pp.476-80. Coleman, David, and Brian Fry. 1991. Carbon Isotope Techniques. Academic Press/Harcourt Brace Jovanovich, New York. De Niro, M. J., and S. Epstein. 1978. Influence of diet on the distribution of carbon isotopes in animals. Geochim. et Cosmochim. Acta v.42 pp.495-506. De Niro, Michael J. 1987. Stable Isotopy and Archaeology. American Scientist v.75 pp.182-187. Ehleringer, J. R. and P.W. Rundel. 1989. Stable Isotopes: History, Units, and Instrumentation. In Rundel, P. W., J. R. Ehleringer and K. A. Nagy, eds. Stable Isotopes in Ecological Research. Springer Verlag, New York. Hall, R. L. 1967. Those late corn dates: Isotopic fractionation as a source of error in carbon-14 dates. Mich. Archaeol. v.13 pp.171-79. Hatch, M. D. and C. R. Slack. 1970. The C4 carboxylic acid pathway of photosynthesis. pp.35-106. In Reinhold L and Liwschitz Y, eds., Progress in Phytochemistry. Wileylnterscience, New York. Hatch, M. D., and C. R. Slack. 1966. Photosynthesis by sugarcane leaves. A new carboxylation reaction and the pathway of sugar formation. Biochem. J. v.101 pp.10311. Hayes, J. M. 1982. Fractionation et. al.: An introduction to isotopic measurements and terminology. Spectra v.8 pp.3-8. Rounick, J. S. and M. J. Winterbourn. 1986. Stable carbon isotopes and carbon flow in ecosystems. BioScience v.36 pp.171-176. Rundel, P. W., J. R. Ehleringer and K. A. Nagy (eds.) 1989. Stable Isotopes in Ecological Research. Springer Verlag, New York. Smith, B. N. and S. Epstein. 1971. Two categories of 13C/12C ratios for higher plants. Plant Physiol. v.47 pp.380-384.
5
van der Merwe, Nikolaas J. 1982. Carbon isotopes, Photosynthesis, and Archaeology. American Scientist, Nov-Dec 1982, 596-606. van der Merwe, N. J., and J. C. Vogel. 1978. 13C content of human collagen as a measure of prehistoric diet in Woodland North America. Nature v.276 pp.815-16. van der Merwe, Nikolaas J. 1982. Carbon isotopes, photosynthesis, and archaeology. American Scientist v.70 pp.596-605. Vogel, J. C. 1978b. Isotopic assessment of the dietary habits of ungulates. South Afr. J. Sci. v.74 pp.298-301. Vogel, J. C. and N. J. van der Merwe. 1977. Isotopic evidence for early maize cultivation in New York State. Am. Antiquity v.42 pp.238-42. Zelitch, I. 1971. Photosynthesis, Photorespiration, and Plant Productivity. Academic Press, New York.
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