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Comment on “Oxygen and carbon isotopic composition of Ordovician brachiopods: Implications for coeval seawater” by H. Qing and J. Veizer
Geochimicaet CosmochimicaActa,Vol. 59, No. 13, pp. 2843-2844, 1995 Copyright© 1995 ElsevierScienceLtd Printed in the USA.All rightsreserved 0016-7037/95 $9.50 + .00
Pergamon
0016-7037(95)00176-X
Comment
Comment on "Oxygen and carbon isotopic composition of Ordovician brachiopods: Implications for coeval seawater" by H. Qing and J. Veizer LYNTON S. LAND Depamnent of Geological Sciences, University of Texas, Austin, "IX 78713, USA (Received February 23, 1995; accepted in revised forra April 14, 1995 ) a lack of alteration. Luminescence is not a requisite product of diagenesis because diagenesis can take place in a system starved for Mn/Fe. The trace elements analyzed by Qing and Veizer (1994) show excellent covariation (e,g., their Fig. 7), a characteristic of mixing between carbonate phases, and not of primary skeletal carbonate. The oxygen isotopic composition of carbonates is the parameter that is normally measured which is most sensitive to alteration, and any detectable alteration in trace element composition (613C or 87sr/a6sr) proves that the 6 ~80 is not primary, although by happenstance, it may not have changed (Land, 1986; Banner and Hanson, 1990). Several authors (Rush and Chafetz, 1990; Banner and Kaufman, 1994) have demonstrated conclusively that despite careful sampling, the criteria "no luminescence, low Mn, and high Sr" are flawed, and do not preclude altered 8 ~80 values. The inability of Qing and Veizer (1994) to explain "anomalous" -8%o data, which otherwise pass their "quality control" (p. 4433) is very simply explained--everything is altered ! Microscopic examination of cathodoluminescence does not even closely approach the scale of the chemical bonds in the crystals being analyzed. Chemical bonds must remain intact for hundreds of millions of years if the 8180 of the crystals is to be preserved. The statement of Qing and Veizer's (1994) that "Only a substantial dissolution/reprecipitation, usually detectable by SEM, can result in large-scale resetting of the original isotopic signal." (p. 4432) is merely megascopic wishful thinking when considered at the scale of the crystal lattice. Any technique above the scale of observation in the TEM is too gross to assess the insidious nature of replacement reactions. Comparisons between modem samples and "unaltered" ancient specimens at the TEM level are badly needed to try to assess whether or not replacement might have occurred in the organic matter-rich skeletons which have been analyzed. It is most unlikely that the 6 lsO of the epeiric Ordovician ocean varied by 3%0 (or its temperature varied by 13°C) at times when normal marine brachiopods flourished, as suggested by "best preserved" samples of Qing and Veizer (1994; their Fig. 2). The most ~sO-depleted of their "best preserved" samples are certainly altered. Gao and Land ( 1991 ) documented microquartz chert from Ordovician platform limestones as heavy as +28.5%0 (SMOW), which would be consistent with formation from
On the basis of analyses of brachiopods, Qing and Veizer (1994) proposed that "The observed 6~80 t r e n d . . . (from a minimum of - 10.5%o in the earliest Ordovician to a maximum of -1.5%o at the Ordovician/Silurian boundary) . . . may reflect a progressive cooling during Ordovician, perhaps complemented by a changing ~5~so values of seawater." The data are more likely the result of dlagenesis, providing no information about the 6 lsO and/or temperature of the Ordovician ocean. Gao and Land ( 1991 ) analyzed approximately 120 samples of homogeneous micritic limestone from lower Ordovician strata in Oklahoma, USA. Where the age of the data-sets overlap ( e.g., at 480 Ma), the 61so values of micrite are essentially identical to the ~5'sO values of the brachiopods analyzed by Qing and Veizer (1994). There is no doubt that the ~5~80 of the micrite analyzed by Gao and Land ( 1991 ) records no information whatsoever about the temperature/6 ~80 of Ordovician seawater. The original mineralogy of micritic Ordovician sediment is not known, but Ordovician lime mud certainly did not consist of the trace element-depleted, nonporous mosaics of calcite crystals which constitute the rocks today. Whatever the primary mineralogy of the sediment, it consisted of unconsolidated grains subject to processes such as bioturbation, desiccation, and current lamination exactly like modem sediments. Retention of the oxygen isotopic composition of a primary lime mud during lithification can only happen if mineral stabilization and cementation takes place in a completely open system at surficial temperatures in seawater. This scenario is most unlikely. Subsequent to sedimentation, meteoric diagenesis (resulting from subareal exposure) and burial diagenesis both irrevocably alter the dit80 of the sediment toward depleted values as the sediment is converted into a rock. Even burial diagenesis in a closed but compacting system produces 18t-depleted rocks (Land, 1980). The Ordovician micrites analyzed by Gao and Land (1991) apparently did acquire 18t-depleted values as mineralogic stabilization and cementation progressed. The fact that the brachiopods analyzed by Qing and Veizer (1994) have essentially the same/5 ~80 as the micrites suggests that the brachiopods have also been altered. Qing and Veizer (1994) applied a " standard" set of criteria (no luminescence, low Mn, and high Sr) to test for alteration (e.g., Carpenter and Lohmann, 1989). Lack of luminescence only signifies a lack of the correct proportions of Mn/Fe, not 2843
2844
L.S. Land
-4%o water at 25°C. In the unlikely event that the chert present in the rocks today accurately records the conditions of initial emplacement, then the Ordovician ocean could not have been more depleted than -4%0. Abundant evidence of subareal exposure exists in the rocks Gao and Land (1991) studied. These rocks, like those studied by Qing and Veizer (1994), were undoubtedly affected by meteoric water, probably many times, early in their burial history. The presence of relict lepisphere textures in the chert demands that, like the micrite, the chert today is not the primary sediment but an alteration of the primary sediment. Alteration of a primary (?) opal-CT lepisphere-bearing precursor chert to microquartz chert in meteoric water is the preferred explanation of Gao and Land (1991). Coastal meteoric water having a 6180 of -4%0 is inconsistent with the 1SO-depleted ocean proposed by Qing and Veizer (1994). The alternative, that the ocean was not ~80-depleted, and that the micritic limestones, the brachiopods, and the chert have all been altered, is more in keeping with the data. There is a lementable, but growing tendency in the literature toward acceptance of 6~80 data as being indicative of unaltered marine compositions, as long as they are obtained using the kinds of arbitrary criteria used by Qing and Veizer (1994) (e.g., Carpenter et al., 1991). More data of the sort presented by Qing and Veizer (1994) which purport to document secular 180 variations in the ancient ocean do not make it so. Plausible data apparently can be obtained where skeletal material is extremely well preserved, especially if large specimens are entombed in "impermeable" shales, and when isotopic studies are regional in nature, and are coupled with careful taxonomy, facies interpretation, and paleogeography (e.g., Grossman et al., 1993). These data document an ~so-invarient ocean. Of the three options which seek to explain why sedimentary rocks are depleted in lsO with increasing age, replacement during diagenesis remains the best choice. Grossly altered Earth temperatures are contested by the evolution of a normal marine brachiopod fauna. Data from ophiolites (Gregory, 1991 ), and our current understanding of ocean floor processes (Muehlenbachs and Clayton, 1976; Holland, 1984) are inconsistent with wild, rapid excursions in • 60water~8 Urey et al. (1961) hope for accurate determination of ancient surficial paleotemperatures is yet to be achieved for rocks as old as these, despite the proliferation of mass-spectrometers fostered by their classic research.
REFERENCES
Banner J. L. and Hanson G. N. (1990) Calculation of simultaneous trace element and isotopic variations during water-rock interaction with applications to carbonate diagenesis. Geochim. Cosmochim. Acta 54, 3123-3137. Banner J. L. and Kaufman J. (1994) The isotope record of ocean chemistry and diagenesis preserved in non-luminescent brachiopods from Mississippian carbonate rocks, Illinois and Missouri. GSA Bull. 106, 1074-1082. Carpenter S. J. and Lohmann K. C. (1989) 613C and 6180 variations in late Devonian marine cements from the Golden Spike and Nevis reefs, Alberta, Canada. J. Sediment. Petrol. 59, 792-814. Carpenter S. J., Lohmann K. C., Holden P., Walter L. M., Huston T. J., and Halliday A. N. ( 1991 ) 6 IsO values, 87Sr/~6Srand Sr/Mg ratios of Late Devonian abiotic marine calcite: implications for the composition of ancient seawater. Geochim. Cosmochim. Acta 55, 1991-2010. Gao G. and Land L. S. ( 1991 ) Geochemistry of Cambro-Ordovician Arbuckle limestone, Oklahoma: Implications for diagenetic 6180 alteration and secular 613C and 87Sr/arsr variation. Geochim. Cosmochim. Acta 55, 2911-2920. Gao G. and Land L. S. (1991) Nodular chert from the Arbuekle Group, Slick Hills, SW Oklahoma: A combined field, petrographic and isotopic study. Sedimentology 38, 857-870. Gregory R. T. ( 1991 ) Oxygen isotope history of seawater revisited: Time scales for boundary event changes in the oxygen isotope composition of seawater. In Stable Isotope Geochemistry: A Tribute to Samuel Epstein (ed. H. P. Taylor Jr. et al.); Geochem. Soc. Spec. Pub 3, 65-76. Grossman E. L., Mii H.-S., and Yancey T. E. (1993) Stable isotopes in late Pennsylvanianbrachiopods from the United States: Implications for Carboniferous paleogeography. GSA Bull. 105, 12841296. Holland H. D. (1984) The Chemical Evolution of the Atmosphere and Oceans. Princeton Univ. Press. Land L. S. (1980) The isotopic and trace element geochemistry of dolomite: The state of the art. In Concepts and Models of Dolomitization (ed. D. H. Zenger et al.); SEPM Special Publication 28, 87-110. Land L. S. (1986) Limestone diagenesis--some geochemical considerations. In Studies in Diagenesis (ed. F. A. Mumpton); USGS Bull. 1578, 129-137. Muehlenbachs K. and Clayton R. N. (1976) Oxygen isotope composition of the ocean crust and its bearing on seawater. J. Geophys. Res. 821, 4365-4369. Qing H. and Veizer J. (1994) Oxygen and carbon isotopic composition of Ordovician brachiopods: Implications for coeval seawater. Geochim. Cosmochim. Acta 58, 4429-4442. Rush P. F. and Chafetz H. S. (1990) Fabric-retentive, non-luminescent brachiopods as indicators of original 613C and 6 ~sO composition: A test. J. Sediment. Petrol. 60, 968-981. Urey H. C., Lowenstam H. A., Epstein S., and McKinney C. R. ( 1961) Measurement of paleotemperatures and temperatures of the upper Cretaceous of England, Denmark, and the southeastern United States. GSA Bull. 62, 399-416,
Pergamon
0016-7037(95)00176-X
Comment
Comment on "Oxygen and carbon isotopic composition of Ordovician brachiopods: Implications for coeval seawater" by H. Qing and J. Veizer LYNTON S. LAND Depamnent of Geological Sciences, University of Texas, Austin, "IX 78713, USA (Received February 23, 1995; accepted in revised forra April 14, 1995 ) a lack of alteration. Luminescence is not a requisite product of diagenesis because diagenesis can take place in a system starved for Mn/Fe. The trace elements analyzed by Qing and Veizer (1994) show excellent covariation (e,g., their Fig. 7), a characteristic of mixing between carbonate phases, and not of primary skeletal carbonate. The oxygen isotopic composition of carbonates is the parameter that is normally measured which is most sensitive to alteration, and any detectable alteration in trace element composition (613C or 87sr/a6sr) proves that the 6 ~80 is not primary, although by happenstance, it may not have changed (Land, 1986; Banner and Hanson, 1990). Several authors (Rush and Chafetz, 1990; Banner and Kaufman, 1994) have demonstrated conclusively that despite careful sampling, the criteria "no luminescence, low Mn, and high Sr" are flawed, and do not preclude altered 8 ~80 values. The inability of Qing and Veizer (1994) to explain "anomalous" -8%o data, which otherwise pass their "quality control" (p. 4433) is very simply explained--everything is altered ! Microscopic examination of cathodoluminescence does not even closely approach the scale of the chemical bonds in the crystals being analyzed. Chemical bonds must remain intact for hundreds of millions of years if the 8180 of the crystals is to be preserved. The statement of Qing and Veizer's (1994) that "Only a substantial dissolution/reprecipitation, usually detectable by SEM, can result in large-scale resetting of the original isotopic signal." (p. 4432) is merely megascopic wishful thinking when considered at the scale of the crystal lattice. Any technique above the scale of observation in the TEM is too gross to assess the insidious nature of replacement reactions. Comparisons between modem samples and "unaltered" ancient specimens at the TEM level are badly needed to try to assess whether or not replacement might have occurred in the organic matter-rich skeletons which have been analyzed. It is most unlikely that the 6 lsO of the epeiric Ordovician ocean varied by 3%0 (or its temperature varied by 13°C) at times when normal marine brachiopods flourished, as suggested by "best preserved" samples of Qing and Veizer (1994; their Fig. 2). The most ~sO-depleted of their "best preserved" samples are certainly altered. Gao and Land ( 1991 ) documented microquartz chert from Ordovician platform limestones as heavy as +28.5%0 (SMOW), which would be consistent with formation from
On the basis of analyses of brachiopods, Qing and Veizer (1994) proposed that "The observed 6~80 t r e n d . . . (from a minimum of - 10.5%o in the earliest Ordovician to a maximum of -1.5%o at the Ordovician/Silurian boundary) . . . may reflect a progressive cooling during Ordovician, perhaps complemented by a changing ~5~so values of seawater." The data are more likely the result of dlagenesis, providing no information about the 6 lsO and/or temperature of the Ordovician ocean. Gao and Land ( 1991 ) analyzed approximately 120 samples of homogeneous micritic limestone from lower Ordovician strata in Oklahoma, USA. Where the age of the data-sets overlap ( e.g., at 480 Ma), the 61so values of micrite are essentially identical to the ~5'sO values of the brachiopods analyzed by Qing and Veizer (1994). There is no doubt that the ~5~80 of the micrite analyzed by Gao and Land ( 1991 ) records no information whatsoever about the temperature/6 ~80 of Ordovician seawater. The original mineralogy of micritic Ordovician sediment is not known, but Ordovician lime mud certainly did not consist of the trace element-depleted, nonporous mosaics of calcite crystals which constitute the rocks today. Whatever the primary mineralogy of the sediment, it consisted of unconsolidated grains subject to processes such as bioturbation, desiccation, and current lamination exactly like modem sediments. Retention of the oxygen isotopic composition of a primary lime mud during lithification can only happen if mineral stabilization and cementation takes place in a completely open system at surficial temperatures in seawater. This scenario is most unlikely. Subsequent to sedimentation, meteoric diagenesis (resulting from subareal exposure) and burial diagenesis both irrevocably alter the dit80 of the sediment toward depleted values as the sediment is converted into a rock. Even burial diagenesis in a closed but compacting system produces 18t-depleted rocks (Land, 1980). The Ordovician micrites analyzed by Gao and Land (1991) apparently did acquire 18t-depleted values as mineralogic stabilization and cementation progressed. The fact that the brachiopods analyzed by Qing and Veizer (1994) have essentially the same/5 ~80 as the micrites suggests that the brachiopods have also been altered. Qing and Veizer (1994) applied a " standard" set of criteria (no luminescence, low Mn, and high Sr) to test for alteration (e.g., Carpenter and Lohmann, 1989). Lack of luminescence only signifies a lack of the correct proportions of Mn/Fe, not 2843
2844
L.S. Land
-4%o water at 25°C. In the unlikely event that the chert present in the rocks today accurately records the conditions of initial emplacement, then the Ordovician ocean could not have been more depleted than -4%0. Abundant evidence of subareal exposure exists in the rocks Gao and Land (1991) studied. These rocks, like those studied by Qing and Veizer (1994), were undoubtedly affected by meteoric water, probably many times, early in their burial history. The presence of relict lepisphere textures in the chert demands that, like the micrite, the chert today is not the primary sediment but an alteration of the primary sediment. Alteration of a primary (?) opal-CT lepisphere-bearing precursor chert to microquartz chert in meteoric water is the preferred explanation of Gao and Land (1991). Coastal meteoric water having a 6180 of -4%0 is inconsistent with the 1SO-depleted ocean proposed by Qing and Veizer (1994). The alternative, that the ocean was not ~80-depleted, and that the micritic limestones, the brachiopods, and the chert have all been altered, is more in keeping with the data. There is a lementable, but growing tendency in the literature toward acceptance of 6~80 data as being indicative of unaltered marine compositions, as long as they are obtained using the kinds of arbitrary criteria used by Qing and Veizer (1994) (e.g., Carpenter et al., 1991). More data of the sort presented by Qing and Veizer (1994) which purport to document secular 180 variations in the ancient ocean do not make it so. Plausible data apparently can be obtained where skeletal material is extremely well preserved, especially if large specimens are entombed in "impermeable" shales, and when isotopic studies are regional in nature, and are coupled with careful taxonomy, facies interpretation, and paleogeography (e.g., Grossman et al., 1993). These data document an ~so-invarient ocean. Of the three options which seek to explain why sedimentary rocks are depleted in lsO with increasing age, replacement during diagenesis remains the best choice. Grossly altered Earth temperatures are contested by the evolution of a normal marine brachiopod fauna. Data from ophiolites (Gregory, 1991 ), and our current understanding of ocean floor processes (Muehlenbachs and Clayton, 1976; Holland, 1984) are inconsistent with wild, rapid excursions in • 60water~8 Urey et al. (1961) hope for accurate determination of ancient surficial paleotemperatures is yet to be achieved for rocks as old as these, despite the proliferation of mass-spectrometers fostered by their classic research.
REFERENCES
Banner J. L. and Hanson G. N. (1990) Calculation of simultaneous trace element and isotopic variations during water-rock interaction with applications to carbonate diagenesis. Geochim. Cosmochim. Acta 54, 3123-3137. Banner J. L. and Kaufman J. (1994) The isotope record of ocean chemistry and diagenesis preserved in non-luminescent brachiopods from Mississippian carbonate rocks, Illinois and Missouri. GSA Bull. 106, 1074-1082. Carpenter S. J. and Lohmann K. C. (1989) 613C and 6180 variations in late Devonian marine cements from the Golden Spike and Nevis reefs, Alberta, Canada. J. Sediment. Petrol. 59, 792-814. Carpenter S. J., Lohmann K. C., Holden P., Walter L. M., Huston T. J., and Halliday A. N. ( 1991 ) 6 IsO values, 87Sr/~6Srand Sr/Mg ratios of Late Devonian abiotic marine calcite: implications for the composition of ancient seawater. Geochim. Cosmochim. Acta 55, 1991-2010. Gao G. and Land L. S. ( 1991 ) Geochemistry of Cambro-Ordovician Arbuckle limestone, Oklahoma: Implications for diagenetic 6180 alteration and secular 613C and 87Sr/arsr variation. Geochim. Cosmochim. Acta 55, 2911-2920. Gao G. and Land L. S. (1991) Nodular chert from the Arbuekle Group, Slick Hills, SW Oklahoma: A combined field, petrographic and isotopic study. Sedimentology 38, 857-870. Gregory R. T. ( 1991 ) Oxygen isotope history of seawater revisited: Time scales for boundary event changes in the oxygen isotope composition of seawater. In Stable Isotope Geochemistry: A Tribute to Samuel Epstein (ed. H. P. Taylor Jr. et al.); Geochem. Soc. Spec. Pub 3, 65-76. Grossman E. L., Mii H.-S., and Yancey T. E. (1993) Stable isotopes in late Pennsylvanianbrachiopods from the United States: Implications for Carboniferous paleogeography. GSA Bull. 105, 12841296. Holland H. D. (1984) The Chemical Evolution of the Atmosphere and Oceans. Princeton Univ. Press. Land L. S. (1980) The isotopic and trace element geochemistry of dolomite: The state of the art. In Concepts and Models of Dolomitization (ed. D. H. Zenger et al.); SEPM Special Publication 28, 87-110. Land L. S. (1986) Limestone diagenesis--some geochemical considerations. In Studies in Diagenesis (ed. F. A. Mumpton); USGS Bull. 1578, 129-137. Muehlenbachs K. and Clayton R. N. (1976) Oxygen isotope composition of the ocean crust and its bearing on seawater. J. Geophys. Res. 821, 4365-4369. Qing H. and Veizer J. (1994) Oxygen and carbon isotopic composition of Ordovician brachiopods: Implications for coeval seawater. Geochim. Cosmochim. Acta 58, 4429-4442. Rush P. F. and Chafetz H. S. (1990) Fabric-retentive, non-luminescent brachiopods as indicators of original 613C and 6 ~sO composition: A test. J. Sediment. Petrol. 60, 968-981. Urey H. C., Lowenstam H. A., Epstein S., and McKinney C. R. ( 1961) Measurement of paleotemperatures and temperatures of the upper Cretaceous of England, Denmark, and the southeastern United States. GSA Bull. 62, 399-416,