applying oxygen isotope paleothermometry in deep time ethan l ...

APPLYING OXYGEN ISOTOPE PALEOTHERMOMETRY IN DEEP TIME 

ETHAN L. GROSSMAN 

Department of Geology and Geophysics, Texas A&M University, 

College Station, TX 77843-3115 USA 

e-grossman@tamu.edu 

ABSTRACT.—Oxygen isotope paleotemperature studies of the Mesozoic and Paleozoic are based mainly 

on conodonts, belemnite guards, and brachiopod shells—material resistant to diagenesis and generally 

precipitated in oxygen isotope equilibrium with ambient water. The greatest obstacle to accurate oxygen 

isotope paleothermometry in deep time is uncertainty in the oxygen isotopic composition of the ambient 

seawater. The second greatest obstacle is fossil diagenesis. Useful application of the oxygen isotope 

method to brachiopod shells requires extreme care in sample screening and analyses, and is best done 

with scanning-electron microscopy, and petrographic and cathodoluminescence microscopy , and traceelement 

analysis. Correct interpretation of oxygen isotope data is greatly aided by thorough understanding 

of the paleolatitude, paleoecology, and depositional environment of the samples. The oxygen isotope record 

for the Triassic, based on brachiopod shells, is too sparse to show any distinct isotopic features. Jurassic 

and Early Cretaceous δ 18 O records, based on belemnites, show a Toarcian (Jurassic) decline (warming), 

a Callovian-Oxfordian acme, and an Early Cretaceous increase (cooling) to a Valanginian- 

Hauterivian maximum, followed by a decline (warming) to a middle Barremian minimum. Deep-time 

applications to oxygen isotope thermometry provide evidence for cooling and glaciation in the Ordovician, 

Carboniferous, and Permian. The δ 18 O values from Silurian and Devonian brachiopod shells and 

conodonts average lower than those of the remaining Phanerozoic because of the absence of continental 

glaciers and possibly higher temperatures (~37°C), although slightly lower (≤2‰) seawater δ 18 O cannot 

be ruled out. The hypothesis of high temperatures in the early Paleozoic implies a relatively constant hydrospheric 

δ 18 O, which is supported by clumped isotope paleotemperatures. However, more research is 

needed to develop methods for evaluating clumped isotope reordering in fossils. Ongoing and future research 

in oxygen isotope and clumped isotope thermometry hold the promise of resolving deep-time temperatures, 

seawater δ 18 O, and salinity with heretofore unavailable accuracy (±2°C, ±0.4‰, and ±2 psu), 

providing the environmental setting for the evolution of metazoan life on Earth. 

INTRODUCTION 

OXYGEN ISOTOPE paleothermometry’s earliest 

application was a deep-time study of Late Cretaceous 

paleotemperatures (Urey et al., 1951). All 

of the concerns raised in Urey et al. (1951) are 

still relevant today—disequilibrium precipitation 

of biogenic carbonate, the constancy of seawater 

δ 18 O, ecologic influences, spatial variability versus 

temporal trends, and the preservation of the 

record through geologic time. This chapter covers 

relations for oxygen isotopic equilibrium in carbonate 

and phosphate mineral, methods to test for 

diagenesis of biogenic carbonates, evidence for 

relative constancy of seawater δ 18 O, examples of 

regional overprinting of the global signal, and 

finally, global compilations for the Mesozoic and 

Paleozoic based mainly on well-preserved 

brachiopod shells, belemnite guards, and conodonts. 

Principles of oxygen isotope thermometry 

Oxygen isotope paleothermometry is founded 

on the temperature dependence of oxygen isotope 

fractionation between authigenic minerals and 

ambient waters. Under equilibrium conditions, the 

18 

O/ 16 O of sedimentary carbonate and phosphate 

minerals depends only on the temperature of precipitation 

and the 18 O/ 16 O of the ambient water. 

Thermodynamic relationships and bond vibrational 

frequencies can be used to determine the 

mineral-water isotopic fractionation relations, but 

not with the precision and accuracy necessary for 

paleothermometry. Such an application requires 

In Reconstructing Earth’s Deep-Time Climate—The State of the Art in 2012, Paleontological Society Short Course, 

November 3, 2012. The Paleontological Society Papers, Volume 18, Linda C. Ivany and Brian T. Huber (eds.), 

pp. 39–67. Copyright © 2012 The Paleontological Society.

THE PALEONTOLOGICAL SOCIETY PAPERS, VOL. 18 

calibrations based on mineral-water oxygen exchange 

experiments at high temperatures, mineral 

precipitation experiments at low temperatures, 

and/or natural experiments using minerals grown 

under known conditions (see also Anderson and 

Arthur, 1983; O’Neil, 1986; Pearson, this volume, 

for reviews). 

Terminology and standardization.— 

Equilibrium isotopic fractionation between minerals 

and water is described by the fractionation 

factor (α) 

where R is the isotopic ratio (e.g., 18 O/ 16 O) and A 

and B are the mineral and water respectively. Isotopic 

ratios are reported in delta (δ) notation relative 

to an internationally accepted standard. The 

equation is defined as 

where Rx and Rstd refer to the 18 O/ 16 O of the 

sample and standard, respectively, and ‰ is per 

mil (parts-per-thousand). For oxygen isotopes, δ 

notation is defined by the equation 

The accepted standards for reporting oxygen isotope 

data are SMOW (Standard Mean Ocean Water) 

and PDB (Peedee Belemnite). SMOW started 

as a hypothetical water approximating the average 

oxygen and hydrogen isotopic composition of 

seawater and defined relative to the NBS1 water 

standard (Craig, 1961). A water standard was later 

mixed by Harmon Craig at Scripps Institution of 

Oceanography to reproduce the hypothetical values 

(Gröning, 2004), becoming the SMOW standard. 

Because the supply of SMOW has been exhausted, 

a new water standard, VSMOW (Vienna 

SMOW), which is analytically indistinguishable 

in δ 18 O from SMOW, was prepared and distributed 

(Gonfiantini, 1984; Gröning, 2004). To 

minimize confusion, the International Atomic Energy 

Agency decided to refer to Craig’s original 

SMOW standard as VSMOW, a convention followed 

in this chapter. Phosphate δ 18 O data are 

also reported versus VSMOW. To account for 

inter-laboratory differences, researchers report the 

value obtained for the phosphorite rock standard 

NBS120 (either aliquot b or c; Vennemann et al., 

2002; Pucéat et al., 2010; MacLeod, this volume). 

The necessity to report NBS120 values is underscored 

by the 0.9‰ range in published values 

(21.7‰, Lécuyer et al., 1996, Trotter et al., 2008; 

22.4‰, Joachimski et al., 2006, Data Repository 

item 2006058; 22.6‰, Vennemann et al., 2002). 

Oxygen isotope data for carbonate minerals are 

usually reported relative to PDB or VPDB. PDB, 

calcite from the belemnite Belemnitella americana 

from the Cretaceous Peedee formation in 

South Carolina, was used as the working standard 

in the pioneering studies at the University of Chicago 

(Urey et al., 1951). PDB powder, however, 

has long been exhausted, so the secondary standards 

NBS-19 and NBS-20 have been used for 

calibration to PDB. For carbonates, the recommended 

practice is calibration to PDB using the 

NBS-19 calcite standard (δ 18 O = -2.20‰ versus 

PDB; Gonfiantini, 1984), referred to as calibration 

to VPDB (Vienna PDB; Coplen et al., 1996). 

The following equations are used to convert data 

between VPDB and VSMOW standardization 

(Coplen, 1988): 

δ 18 Ox-VSMOW = 1.03091 δ 18 Ox-VPDB + 30.91 

and δ 18 Ox-VPDB = 0.970017 δ 18 Ox-VSMOW – 29.98 

Oxygen isotope paleothermometers.—Table 1 

summarizes the commonly used fractionation relations 

and paleotemperature equations, the temperature 

range for which they were determined, 

and the material analyzed. McCrea (1950) and 

Epstein et al. (1951) experimentally demonstrated 

the temperature dependence of oxygen isotope 

fractionation between calcium carbonate and water 

first theorized by Urey (1947), and developed 

preliminary oxygen isotope paleotemperature 

equations based on inorganic and biogenic carbonates 

respectively. Epstein et al. (1953) developed 

the first practical oxygen isotopic paleotemperature 

scale, based on oxygen isotopic measurements 

of biogenic carbonate (mostly mollusk 

shells) and environmental waters (Table 1, Figure 

1). Their equation with an added correction for 

the isotopic composition of the water (Epstein and 

Lowenstam, 1953; Epstein and Mayeda, 1953) is: 

T (°C) = 16.5 – 4.3 (δ 18 OCaCO3– δ 18 Ow-PDB) + 0.14 

(δ 18 OCaCO3– δ 18 Ow-PDB) 2 

40

GROSSMAN: OXYGEN ISOTOPE PALEOTHERMOMETRY IN DEEP TIME 

Temperature (°C) 

5 

35 

4 

3 

2 

30 

1 

0 !"#$% 

-1 

-2 

25 

-3 

-4 

-5 

20 

2 2.5 3 3.5 4 

δ 18 O CaCO3 -δ w-VSMOW (‰) 

15 

10 

Friedman & O'Neil (1977) 

!"#$%&'() 

Kim and O'Neil (1997) 

5 

Erez and Luz (1983) 

Epstein et al. (1953) 

0 

Grossman and Ku (1986, arag.) 

Kim et al. (2007; arag.) 

-5 

-4 -3 -2 -1 0 1 2 3 4 

δ 18 O CaCO3 -δ w-VSMOW (‰) 

FIGURE 1.—Comparison of paleotemperature equations 

for calcite-water and aragonite-water fractionation. 

The inset shows the divergence of paleotemperature 

equations at low temperatures. 

where the δ 18 OCaCO3 and δ 18 Ow-PDB are the δ 18 O of 

calcium carbonate and water relative to PDB. The 

water standardization to “PDB” is a source of 

confusion in the literature. The δ 18 O of H2O oxygen 

was not compared with the δ 18 O of PDB oxygen, 

but instead the δ 18 O of CO2 equilibrated with 

the water sample at 25.3°C was compared with 

CO2 derived from PDB reacted with phosphoric 

acid at 25°C. Since the mass spectrometer measured 

CO2, it was easy to think of the standard in 

terms of the gas rather than the original solid or 

liquid. This standardization yields δ 18 O values for 

water that at first were reported as 0.20‰ lower 

than values reported relative to VSMOW (Craig, 

1965). Later this calibration was updated to 

0.22‰ (Friedman and O’Neil, 1977) and then 

0.27‰ (Hut, 1987). If and what value of the 

“PDB”-VSMOW correction should be applied 

depends on the study (Bemis et al., 1999; Pearson, 

this volume). When using the Epstein et al. (1953) 

equation, one should use the most up-to-date 

“PDB”-VSMOW correction (-0.27‰) as the Epstein 

et al. water values were directly standardized 

with PDB-derived CO2 (contrasts with the use of 

0.20‰ suggested by Bemis et al., 1998). For 

other studies that include a water δ 18 O correction 

relative to “PDB,” but measured seawater δ 18 O 

relative to VSMOW, water δ 18 O values versus 

VSMOW must be corrected by subtracting 0.20, 

0.22, or 0.27‰ depending on the value used by 

the reporting authors (Bemis et al., 1999). These 

Temperature (°C) 

corrections are shown in the paleotemperature 

equations listed in Table 1. Note that the equations 

in Table 1 have been factored to remove the 

extra parentheses so that the Epstein et al. (1953) 

equation with the correction for the water δ 18 O 

versus VSMOW (δ 18 Ow): 

T (°C) = 16.5 – 4.3 (δ 18 OCaCO3– (δ 18 Ow – 0.27)) + 

0.14 (δ 18 OCaCO3– (δ 18 Ow – 0.27)) 2 

becomes 

T (°C) = 16.5 – 4.3 (δ 18 OCaCO3– δ 18 Ow + 0.27) + 

0.14 (δ 18 OCaCO3– δ 18 Ow + 0.27) 2 

Equations that were defined using water δ 18 O values 

relative to VSMOW do not require the added 

correction. 

In an attempt to circumvent the complications 

of explaining standardization of waters to “PDB” 

and yet follow the original intent of Sam Epstein 

and his colleagues, I have referred to water standardization 

to “PDB” as standardization to “average 

marine water” (AMW) in Kobashi et al. 

(2004) and Grossman (2012). My approach was 

based on Epstein et al.’s (1953, p. 1324) observation 

that “the O 18 /O 16 ratio of carbon dioxide 

equilibrated with average marine water was found 

to be 0.0‰ relative to our working standard gas.” 

In Grossman (2012), I suggested a single correction 

factor of -0.27‰, but as pointed out by Bemis 

et al. (1998), for some studies a value of 

0.20‰ or 0.22‰ must be used (Table 1). The error 

introduced by using 0.27‰ instead of 0.20‰ 

is only about 0.3°C, within the error of paleotemperature 

equations. Unfortunately, my idea of referring 

to “PDB” standardization for waters as 

standardization to AMW may add yet more confusion. 

Thus, I recommend using the equations in 

Table 1 as a guide, being sure to cite the original 

studies. 

O’Neil et al. (1969) developed the first practical 

relationship for abiotic calcite-water fractionation 

based on laboratory experiments. Revised by 

Friedman and O’Neil (1977), this equation is: 

1000 lnα = [2.78 x 10 6 /T 2 ] - 2.89 

where T is temperature in kelvin. Subsequent 

laboratory experiments by Kim and O’Neil (1997) 

provided a new equation for calcite-water fractionation: 

1000 lnαcalcite-water = 18.03 (10 3 T -1 ) - 32.42 

41


TABLE 1.—Commonly used 18 O fractionation and 18 O paleotemperature equations for CaCO3 and phosphate. In 

paleotemperature equations, CaCO3 δ 18 O data are versus PDB, whereas phosphate data are versus VSMOW. Water 

data in paleotemperature equations are relative to VSMOW (δ 18 Ow). In some equations, a correction of 0.20‰, 

0.22‰, or 0.27‰ is added where authors referenced water values relative to “PDB” (see text for discussion). 

EQUATION 

CaCO3-H2O fractionation relations 

T 

RANGE 

(°C) 

1000 lnα = 2.78 (10 6 T(K) -2 ) – 3.39 0–500 

1000 lnα = 2.78 (10 6 T(K) -2 ) - 2.89 0–500 

MATERIAL 

Calcite-water 

exchange and 

synthetic calcite 

Calcite-water 

exchange and 

synthetic calcite 

1000 lnαcalcite-water = 18.03 (10 3 T(K) -1 ) - 32.42 10–40 Synthetic calcite 

1000 lnαaragonite-water = 17.88 (±0.13) (10 3 T(K) -1 ) – 

31.14 (±0.46) 

Oxygen isotope paleotemperature equations 

0–40 

T (°C) = 16.5 – 4.3 (δ 18 OCaCO3– δ 18 Ow + 0.27) + 

0.14 (δ 18 OCaCO3– δ 18 Ow + 0.27) 2 7.2–29.5 

T (°C) = 16.0 – 4.14 (δ 18 OCaCO3– δ 18 Ow) + 0.13 

(δ 18 OCaCO3– δ 18 Ow) 2 7.2–29.5 

Synthetic aragonite 

Biogenic aragonite 

and calcite 


and calcite 

T (°C) = 16.9 – 4.38 (δ 18 Ocalcite– δ 18 Ow + 0.20) + 

0.10 (δ 18 Ocalcite– δ 18 Ow + 0.20) 2 Synthetic calcite 

T (°C) = 17.04 – 4.34 (δ 18 Ocalcite– δ 18 Ow + 0.20) + 

0.16 (δ 18 Ocalcite– δ 18 Ow + 0.20) 2 4.5–23.3 

T (°C) = 17.0 - 4.52 (δ 18 Ocalcite - δ 18 Ow + 0.22) + 

0.03 (δ 18 Ocalcite - δ 18 Ow + 0.22) 2 14–30 

Calcite from cultured 

Patinopecten 

yessoensis 

Foraminiferal 

calcite 

T (°C) = 15.7 – 4.36 (δ 18 Ocalcite– δ 18 Ow) + 0.12 

(δ 18 Ocalcite– δ 18 Ow) 2 0–60 Synthetic calcite 

T (°C) = 16.5 – 4.80 (δ 18 Ocalcite - δw + 0.27) (low 

light) 

T (°C) = 14.9 – 4.80 (δ 18 Ocalcite - δw + 0.27) (high 

light) 

15–25 

REFER- 

ENCE 

O’Neil et al. 

(1969) 

Friedman 

and O’Neil 

(1977) 

Kim & 

O’Neil 

(1997) 

Kim et al. 

(2007) 

Epstein et al. 

(1953) 

Anderson 

and Arthur 

(1983) 

Shackleton 

(1974) 

Horibe & 

Oba (1972) 

Erez & Luz 

(1983) 

Hays & 

Grossman 

(1991) 

Planktonic foraminiferal 

calcite Bemis et al. 

(Orbulina universa) 

(1998) 

T (°C) = 16.1 - 4.64 (δ 18 Ocalcite - δ 18 Ow + 0.27) + 

0.09 (δ 18 Ocalcite - δ 18 Ow + 0.27) 2 10–40 Synthetic calcite 

Bemis et al. 

(1998) 

T (°C) = 13.7 - 4.54 (δ 18 Ocalcite - δ 18 Ow) + 0.094 

(δ 18 Ocalcite - δ 18 Ow) 2 10–40 Synthetic calcite This chapter 

T (°C) = 16.1 - 4.76 (δ 18 Ocalcite - δ 18 Ow + 0.27) 4.1–25.6 

T (°C) = 20.6 - 4.34 (δ 18 Oaragonite - δ 18 Ow + 0.20) 2.6–22.0 

T (°C) = 19.7 - 4.34 (δ 18 Oaragonite - δ 18 Ow) 2.6–22.0 

Benthic foraminiferal 

calcite 

(Cibicidoides and 

Planulina) 



Lynch- 

Stieglitz et 

al. (1999) 

Grossman & 

Ku (1986) 

Hudson & 

Anderson 

(1989) 

COMMENTS 

High temperature exchange (200- 

500°C) and calcite synthesis (0 

and 25°C) 

Recalculation of O’Neil et al. 

(1969) using revised αCO2-H2O 

(1.0412) 

Term for water correction added 

in Epstein and Lowenstam (1953) 

Revision of Epstein et al. (1953) 

with δ 18 Ow referenced to 

VSMOW 

Quadratic approximation of 

O’Neil et al. (1969) 

Cultured mollusks, Mutsu Bay, 

Japan 

50 – 90% of foraminiferal test 

grown under controlled conditions 

Quadratic approximation of 

O’Neil et al. (1969; with correction 

of Friedman and O’Neil, 

1977) 

δw-VPDB values are obtained by 

subtracting 0.27‰ from δ 18 O 

values reported versus VSMOW 

Quadratic approximation of Kim 

& O’Neil (1997) using the acid 

fractionation factor of Sharma and 

Clayton (1965; 1000lnα = 10.25) 

Quadratic approximation of Kim 

& O’Neil (1997) using Kim & 

O’Neil’s acid fractionation factor 

(1000lnα = 10.44) 

Surface sediments from Little 

Bahama Bank 

Equation 1 

Equation 1 of Grossman and Ku 

(1986) with water δ 18 O values cast 

in terms of VSMOW 

42


Additional carbonate 18 O paleotemperature relations 

have been defined for calcitic foraminifera 

(Erez and Luz, 1983; Bemis et al., 1998; Lynch- 

Stieglitz et al., 1999), aragonitic mollusks and 

foraminifera (Grossman and Ku, 1986), and synthetic 

aragonite (Kim et al., 2007; Fig. 1). In this 

chapter, I use a quadratic approximation of the 

O’Neil et al. (1969) equation by Hays and Grossman 

(1991) (Table 1): 

T (°C) = 15.7 – 4.36 (δ 18 Ocalcite– δ 18 Ow) + 0.12 

(δ 18 Ocalcite– δ 18 Ow) 2 

where δ 18 Ow is the δ 18 O of water versus VSMOW. 

This equation yields warmer paleotemperatures 

than the Kim and O’Neil (1997) equation (+1.8° 

and +3.3°C at 25° and 0°C respectively), and 

agrees better with data for the deep-sea benthic 

foraminifera Uvigerina (Shackleton, 1974) and 

the Epstein et al. (1953) equation. Which equation 

is most accurate for paleotemperature studies in 

general remains uncertain. Researchers tend to 

use the equation derived from material similar to 

their samples (e.g., foraminiferal studies may use 

Erez and Luz, 1983) and the one that gives the 

most accurate temperatures for modern specimens 

(see Pearson, this volume, for an excellent discussion). 

Two relations were commonly applied to 

oxygen isotopic studies of phosphatic materials, 

one based on phosphate within carbonate shells 

(Longinelli and Nuti, 1973) and the other using 

phosphatic teeth and bone (Kolodny et al., 1983; 

Table 1). These relations give similar paleotemperatures. 

However, a new phosphate-water fractionation 

relation by Pucéat et al. (2010) is offset 

from these previous equations by 2‰! The equation 

is: 

T (°C) = 118.7 – 4.22 [(δ 18 Ophosphate + (22.6 - 

δ 18 ONBS120c)) – δ 18 Ow] 

where δ 18 ONBS120c is the value obtained for the 

standard NBS120c and all values are reported 

relative to VSMOW (see MacLeod, this volume). 

Differences between early and recent paleotemperature 

relations may reflect analytical differences 

between laboratories and differences in 

standardization. 

Influence of δ 18 O of environmental waters.— 

The fundamental limitation of the oxygen isotope 

paleothermometer is that the equation has two 

unknowns, temperature and the δ 18 O of the environmental 

water. In certain environments, such as 

high latitudes where glacial meltwater is a significant 

component of surface waters, or in environments 

influenced by large rivers, seawater δ 18 O 

can have a greater control on carbonate δ 18 O than 

temperature (see Pearson, this volume, for additional 

discussion). Furthermore, vapor transport 

from one ocean basin to another can result in significant 

inter-ocean variability (Broecker, 1989). 

The bulk of modern seawater, represented by 

deep water masses, has a relatively narrow δ 18 O 

range from about -0.6‰ for Antarctic Bottom Water 

(AABW) to 0.1‰ for North American Deep 

Water (NADW; Craig and Gordon, 1965; Bigg 

and Rohling, 2000). However, unrestricted, openocean 

surface waters are much more variable, 

ranging from about -0.5‰ in the Southern Ocean 

to 1.4‰ in the dry subtropical high-pressure zone 

←TABLE 1.—Continued. 

EQUATION 

T 

RANGE 

(°C) 

T (°C) = 111.4 - 4.3 (δ 18 Ophosphate - δ 18 Ow) 3.5–27.3 

T (°C) = 113.3 - 4.38 (δ 18 Ophosphate - δ 18 Ow) 3.5–25 

T (°C) = 118.7 - 4.22 [(δ 18 Ophosphate + (22.6 - 

δ 18 ONBS120c) - δ 18 Ow)] 

Effect of Mg content on calcite δ 18 O 

3.5–28 

0.06‰ per mole % MgCO3 25 

0.17 ± 0.02‰ per mole % MgCO3 25 

MATERIAL 

Phosphate in barnacle 

and mollusk 

shells 

Phosphate in fish 

bones and teeth 

Phosphate in fish 

teeth 

Synthetic Mg calcite 

REFER- 

ENCE 

Longinelli 

& Nuti 

(1973) 

Kolodny et 

al. (1983) 

Pucéat et 

al. 2010 

Synthetic Mg calcite 

Tarutani et 

al. (1969) 

Jimenez- 

Lopez et al. 

(2004) 

COMMENTS 

Temperatures from δ 18 OCaCO3. 

Data reported versus VSMOW. 

Temperatures are site averages. 

43


FIGURE 2.—Gridded data set of surface water δ 18 O calculated from the Schmidt et al. (1999) compilation. 

in the North Atlantic Ocean (Figure 2; GEO- 

SECS, 1987; Schmidt et al., 1999; LeGrande and 

Schmidt, 2006). Oxygen isotope variability in 

surface waters of restricted water bodies, such as 

the Arctic Ocean, Mediterranean Sea, and Red 

Sea, is roughly -2‰ to +2‰ (Fig. 2; Schmidt et 

al., 1999; Al-Rousan et al., 2003). To provide a 

first-order correction for the effects of latitudinal 

variation in precipitation minus evaporation (P- 

E), Zachos et al. (1994) developed the following 

empirical relationship for Southern Hemisphere 

surface seawaters (0–70°S) based on GEOSECS 

(1987) data: 

δ 18 Osw (‰ VSMOW) = 0.576 + 0.041x- 0.0017x 2 

+ 1.35·10 -5 x 3 (R = 0.9) 

where x is absolute latitude in degrees. 

Closer to continents, mixing with freshwater 

can lower δ 18 Ow by an amount dependent upon 

the δ 18 Ow of the river water. For example, Mississippi 

River water (δ 18 Ow ≈ -6‰) can lower seawater 

δ 18 O by 0.19‰ per psu (DiMarco et al., 

2012). In the high latitudes of the North Atlantic, 

where freshwater input is more 18 O-depleted 

(-19‰), seawater δ 18 O may decline 0.55‰ per 

psu (LeGrande and Schmidt, 2006). Thus, a 1‰ 

decrease in high-latitude carbonates could reflect 

4–5°C temperature increase or a 2 psu decrease in 

salinity. Because of the lower δ 18 O of highlatitude 

precipitation (Rozanski et al., 1993), shallow 

high-latitude carbonates tend to be more variable 

in δ 18 O than low-latitude carbonates (e.g., 

Tripati et al., 2001; Müller-Lupp and Bauch, 

2005). 

Evaporation of seawater also can have a significant 

effect, with δ 18 O/S slopes of 0.28–0.35‰ 

per psu in the Red Sea and Mediterranean Sea 

(Railsback et al., 1989; LeGrande and Schmidt, 

2006). Complicating the paradigm that freshwater 

input causes δ 18 O depletion in waters and carbonates, 

Lloyd (1964) and later Swart and Price 

(2002) showed that low salinity waters in Florida 

Bay (20–30 psu), sourced by freshwaters from the 

Everglades, can have δ 18 O values 1–3‰ higher 

than those of open ocean waters. Though an extreme 

case, such an environment was possible in 

the subtropical epicontinental seas of Paleozoic 

North America. 

River discharge can have an impact on waters 

far from the river mouth. During the spring, discharge 

from the Mississippi-Atchafalaya River 

system into the Gulf of Mexico lowers surface 

salinities and seawater δ 18 O above Stetson Bank, 

150 km offshore and 450 km from the outlet, by 4 

psu and 0.6‰ respectively (Gentry et al., 2008). 

44


This is especially relevant when considering the 

potential of freshwater input into restricted seas, 

such as the Paleozoic epicontinental seas of North 

America. To minimize the effects of salinity 

variation in paleotemperature studies (unless that 

is the signal of interest), researchers analyze 

stenohaline taxa and consider the paleolatitude 

and paleoaltitude of the catchment area of regional 

discharge. 

Lastly, the carbonate ion concentration (and 

thus pH) of ambient waters has been demonstrated 

to influence planktonic foraminiferal δ 18 O 

values. (Spero et al., 1997). Zeebe (1999) proposed 

that this pH dependence reflects the proportion 

of CaCO3 oxygen derived from bicarbonate 

ion, carbonate ion, and aqueous CO2, each of 

which has a different δ 18 O. For the pH range of 

modern seawater, 7.6 to 8.4, this equates to a δ 18 O 

range of ~1‰, with high δ 18 O at low pH (Beck et 

al., 2005). It is not yet known whether metazoan 

carbonates or phosphates exhibit this pHdependent 

isotope effect. 

Samples and methods 

Introduction.—Fossils used for oxygen isotope 

paleothermometry should be geographically 

and stratigraphically widespread, easy to sample, 

and resistant to diagenesis. Furthermore, fossils 

should be precipitated in 18 O equilibrium with the 

ambient water, or at a constant offset from equilibrium. 

Disequilibrium fractionation in biogenic 

carbonate, termed “vital effect” (Urey et al., 

1951), is closely tied to taxonomy and physiology. 

Mollusks, brachiopods, sclerosponges, and many 

smaller foraminifera typically secrete skeletons at 

or near oxygen isotopic equilibrium (e.g., 

González and Lohmann, 1985; Grossman, 1987; 

Wefer and Berger, 1991; Swart et al., 1998; Figure 

3), whereas corals, echinoderms, and larger 

benthic foraminifera precipitate CaCO3 with δ 18 O 

values as much as 3‰ lower than equilibrium 

values. Taxa exhibiting vital effects, such as corals, 

often precipitate carbonate with a relatively 

constant δ 18 O offset from equilibrium values (e.g., 

Leder et al., 1996). 

Habitat.—The environmental information 

archived in the isotopic composition of fossils 

depends on the organism’s habitat (Figure 4). 

Taxa that are planktonic and depend on photosynthesis 

either directly or indirectly, such as the 

planktonic foraminifer Globigerinoides ruber, 

occupy the photic zone. These taxa provide the 

advantage of a known paleodepth. In contrast, 

benthic taxa, such as brachiopods, bivalves, and 

Aragonite cement 

Mg calcite cement 

Red algae 

Green algae 

Encrusting forams 

Mollusks 

Worm tubes 

Corals 

Internal sediments 

Well cemented 

Poorly cemented 

-5 

-4 

-3 

LMC 

Aragonite 

HMC 

most gastropods, grow on the sea floor, the paleodepth 

of which must be constrained by sedimentological 

and paleoecologic depth indicators (e.g., 

green algae, hermatypic coral). Nektonic species 

such as belemnites and conodonts can have variable 

depth habitats, with even diurnal variations, 

or can be nektobenthic. The paleoecology of these 

organisms can be resolved through δ 18 O comparisons 

with co-occurring benthic and planktonic 

species or other species of the same group 

(Wright, 1987; Anderson et al., 1994; Malchus 

and Steuber, 2002; Moriya et al., 2003; Dutton et 

al., 2007; Ivany, this volume; MacLeod, this volume). 

Diagenesis.—The integrity of samples used in 

deep-time paleotemperature studies is paramount 

and has been the subject of lively debate (e.g., 

Land, 1995; Veizer et al., 1995). The oxygen isotopic 

compositions of fossils are altered through 

oxygen exchange with diagenetic waters. Alteration 

is promoted by the higher temperatures and 

pressures of burial. Low-magnesium calcite is 

more resistant to diagenesis than high magnesium 

calcite or aragonite, and persists longer in the 

sedimentary record (see Marshall, 1992, Veizer, 

-2 

-1 

δ 13 C PDB 

5 

4 

3 

2 

1 

δ 18 O PDB 

FIGURE 3.—Oxygen and carbon isotopic compositions 

of components from a coral boundstone from Enewetak 

Atoll, western tropical Pacific (González and 

Lohmann, 1985). The sample is encrusted with foraminifera 

and red algae and bored by worms and bivalves. 

Note the taxonomic differences in isotopic 

composition attributed to vital effect. Boxes represent 

equilibrium fields calculated by the original authors. 

LMC = low Mg calcite, HMC = high Mg calcite. 

-1 

-2 

1 

45


!"#$%&%$#'#()"*)+*),-.$"*(/)#)&(0* 

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&*1'#!'&+2:$ ;#-$/'$0-''/$ 

*10*'>$8'/+2)($ ,#-*3)/),'-*G:$$ 

FIGURE 4.—Schematic flowchart for habitat considerations in the interpretation of oxygen isotope data from fossils 

and microfossils. 

1992, and Corfield, 1995 for reviews on diagenesis 

and stable isotopic signatures). Under conditions 

that isolate fossils from diagenetic fluids, 

such as encapsulation in asphalt or shale (e.g., 

Pennsylvanian Boggy Formation and Holder 

Formation; Brand, 1982; Dickson et al., 1996; 

Seuss et al., 2012), metastable minerals such as 

aragonite and high-Mg calcite may persist. Unlike 

carbonate ions, phosphate ions do not readily exchange 

their oxygen with water; thus biogenic 

apatite δ 18 O is more resistant to diagenesis than 

low-Mg calcite δ 18 O (e.g., Luz et al., 1984; Wenzel 

et al., 2000; MacLeod, this volume). 

Most studies of oxygen isotope paleothermometry 

for Early Cretaceous and older sediments 

rely on brachiopod shells, belemnite 

guards, and conodonts as geochemical archives. 

The absence of deep-sea floor older than ~180 Ma 

means that recovery of early Mesozoic and Paleozoic 

fossils is restricted to sediments from continental 

margins and epicontinental seas. These 

samples are more likely to be subjected to meteoric 

diagenesis and influenced by coastal processes. 

Furthermore, there is an inherent bias toward 

sediments deposited during high sea levels. 

Because of their relatively large size, calcitic mineralogy, 

and dense microcrystalline structure, belemnites 

were used by Urey et al. (1951) in their 

pioneering study of Cretaceous climate. Besides 

belemnites, researchers of Mesozoic paleotemperatures 

have made use of brachiopod shells 

(Korte et al., 2005a) and rare occurrences of 

aragonitic mollusk shells (e.g., Stahl and Jordan, 

1969; Anderson et al., 1994; Malchus and Steuber, 

2002; Nützel et al., 2010). 

Studies of Paleozoic paleoclimate favor articulate 

brachiopod shells because of their wide 

stratigraphic distribution, time range (Early Cambrian 

to Recent), abundance, and resistance to 

diagenesis (Compston, 1960; Lowenstam, 1961; 

Veizer et al., 1986; Popp et al., 1986; Grossman, 

1994). The resistance to diagenesis results from 

their calcitic mineralogy, low magnesium content, 

relatively large size and thickness, and dense microstructure. 

Because of these qualities, brachiopod 

shells are typically 2–3‰ higher in δ 18 O than 

the encasing, diagenetically-modified bulk carbonate 

(Veizer et al., 1999; Mii et al., 1999). 

Growing utilization of conodonts in Paleozoic 

studies reflects their widespread occurrence, their 

resistance to oxygen isotopic alteration, and improvements 

in analytical technique (e.g., Wenzel 

et al., 2000; Joachimski et al., 2004; MacLeod, 

this volume). 

Sample screening.—The effect of diagenesis 

on mineral δ 18 O depends on depositional and diagenetic 

environment, as well as the original δ 18 O 

of the fossil. Cenozoic planktonic foraminifera 

46


High preservation potential 

Low preservation 

potential 

Thick fossil skeletons, 

low-Mg calcite 

(brachiopods, 

belemnites) 

Aragonite mollusks, 

muds;; high Mg 

calcite 

no 

Are specimens texturally 

d 

well preserved in hand 

sample 

Are specimens texturally 

d 

well preserved in hand 

sample 

yes 

no 

yes 

no 

Are shell microstructure 

and crystals well preserved 

(e.g., absence d of pitting, 

overgrowths, or recrystallization) 

and inclusion-free 

Is aragonite or high Mg 

d 

calcite preserved 

yes 

no 

yes 

Are crystals well 

preserved (no d pitting or 

overgrowths in SEM) 

no 

no 

Nonluminescent d 

yes 

Low Fe 

yes 

Proceed with 

confidence! 

If no, then 

cathodoluminescence 

not a useful indicator. 

yes 

Expected 

Mg/Ca and Sr/Ca for 

taxon and 87 Sr/ 86 Sr for 

age 

yes 

Other tests that build confidence when 

affirmative: 

All co-occurring fossils well-preserved 

Species effects in δ 13 C preserved 

δ 18 O range reasonable for the 

environment 

Proceed with 

caution! 

FIGURE 5.—Schematic flowchart of procedures for screening specimens for diagenesis. For more information, see 

Carpenter et al. (1991), Marshall (1992), Grossman (1994), Sharp (2007), and Cochran et al. (2010). 

growing in warm surface water can recrystallize 

on the cold, deep-ocean floor, altering δ 18 O to 

higher values (Schrag, 1999; Pearson et al., 2001). 

Marine fossils altered by 18 O-depleted meteoric 

waters in outcrop and in the terrestrial subsurface 

typically have lower δ 18 O values, resulting in 

higher apparent isotopic temperatures (e.g., Marshall, 

1992). 

Essential in the application of oxygen isotope 

paleothermometry (or any geochemical proxy 

technique) is the ability to identify chemical alteration 

in samples using criteria independent of 

the isotopic values themselves (Compston, 1960). 

Thus, researchers have established protocols 

based on textural and trace-elemental information 

to screen specimens for alteration (Carpenter et 

al., 1991; Marshall, 1992; Grossman, 1994). Early 

researchers established the use of petrographic 

microscopy of thin sections (Compston, 1960), 

cathodoluminescence petrography (Popp et al., 

1986), and trace element chemistry (Veizer et al., 

1986; Popp et al., 1986). Since no method is foolproof, 

a combination of methods is best. 

The first step in sample screening always is 

evaluation of textural preservation in hand sample 

(Figure 5). Specimens showing extensive silica 

47


FIGURE 6.—Series of six matched plane-transmitted light (PL; left) and cathodoluminescence (right) photomicrographs 

of thin sections of brachiopod shells from Carboniferous sediments in West Virginia and Illinois (USA). Thin 

sections were examined under a petrographic microscope using a TECHNOSYN Model 8200 MKII cathodoluminescence 

stage. The operating conditions were gun current of 200-300 mA and voltage of 10–15 kV. Shells were 

imaged using a Coolsnap-Procf camera attached to a desktop computer. For cathodoluminescence images, exposure 

times were 20 s for more luminescent shells and 60 s for those that needed additional time to enhance contrast. The 

images show a gradational scale of cathodoluminescence (black to bright orange), with sites (white boxes) characterized 

as nonluminescent (NL), slightly luminescent (SL), cathodoluminescent (CL), or some combination. Shown 

below the paired images are brachiopod genus, stratigraphic formation, sample locality, stratigraphic stage (North 

American), and specimen and image number. Poor shell preservation is indicated by opaque areas in planetransmitted 

light (Figures 6A,C,E,G.I,K) and by cathodoluminescence (Figures 6B, D, F, H ,J, L). Note that there 

are no standard practices for exposure time for image capture, electron beam current, or camera type. All these variables 

can influence image intensity and thus cathodoluminescence intensity. Furthermore, translucence of PL images 

is influenced thin-section polish and thickness. These thin-sections are relatively thick (300–400 µm) and thus 

less clear, and are treated with a final polish using 0.3 µm Al oxide grit. Image width is 3.25 mm. Modified from 

Flake 2011; Flake et al., in prep. 

replacement and fracturing should be avoided. 

The next step is evaluation of preservation of 

skeletal microstructure using petrographic microscopy 

and/or scanning electron microscopy 

(SEM). Petrographic microscopy is combined 

with cathodoluminescence microscopy to test for 

chemical alteration. The trace element chemistry 

of biogenic carbonates typically is distinct from 

diagenetic calcite. For example, biogenic calcite 

is poor in Mn and Fe because these trace elements 

are insoluble in oxic waters. In contrast, diagenetic 

waters are commonly anoxic, with dissolved 

Mn 2+ and Fe 2+ that easily substitute into the calcite 

lattice. Thus, modern brachiopods exhibit low 

Mn and Fe contents of generally


et al., 1994). Mn concentrations above ~20 ppm 

in calcite can activate cathodoluminescence, 

while Fe concentrations as low as 35 ppm begin 

to quench it, with CL brightness proportional to 

Mn/Fe ratio (Machel, 1985; Mason, 1987; Savard 

et al., 1995). Shells altered in oxic waters may not 

gain Mn and thus may not luminesce (e.g., Rush 

and Chafatz, 1990; Banner and Kaufman, 1994). 

In such cases, screening must rely on other criteria 

such as Sr, Na, and S contents and textural 

preservation (Veizer et al., 1986; Grossman, 

1994). Such trace element tests are not always 

conclusive, but researchers have found SEM effective 

in identifying recrystallization and cementation 

in brachiopod shells, even when such features 

are not visible in thin section (Wenzel, 

2000). 

Most isotopic studies of Paleozoic brachiopods 

screen specimens using either CL microscopy 

and microsampling, or trace element analyses 

of larger, crushed-shell samples. In the 

“TAMU” (Texas A&M University) method, every 

specimen is thin-sectioned and imaged in transmitted 

PL and CL. Shell areas that are clear in 

transmitted light and nonluminescent (NL) are 

microsampled (50 to 150 µg) from thin-sections 

or complementary billets using a dental pick or 

drill (e.g., Grossman et al., 1991, 1993; Mii et al., 

1999, 2001). Areas that are dark or cloudy (secondary 

inclusions) in transmitted light, show faint 

or dull CL, or have fine-scale CL microfractures 

are avoided. At least three shell areas are analyzed 

per thin-section or billet, as well as one area with 

cement or matrix when possible to provide an indication 

of isotopic sensitivity to diagenesis 

(Grossman, 1994). As seen in Figure 7, the δ 18 O 

and δ 13 C values of CL brachiopod shell typically 

are depleted in heavy isotopes compared with NL 

shell. 

While Mn and Fe are often introduced with 

diagenesis, other trace elements are often lost 

(Veizer, 1983; Mii et al., 1999). For example, Figure 

8 shows depletion of S, Na, and Sr in diagenetically 

altered (CL) shell relative to NL shell. 

These trace-element relations underpin the screening 

methods of Ján Veizer and his students and 

colleagues (“Ruhr” method). They crush shells, 

hand-pick 4–6 mg of calcite fragments with a 

binocular microscope, and isotopically and 

chemically analyze the fragments (Bruckschen 

and Veizer, 1997; Veizer et al., 1999). Bruckschen 

et al. (1999, 2001) analyzed brachiopod shells 

from the Donets Basin (Ukraine) and the Moscow 

Basin using both methods. Initial analyses of the 

FIGURE 7.—Isotopic comparison of nonluminescent 

(NL) and luminescent (CL) calcite from spiriferid 

brachiopod shells from Kansas and New Mexico 

(USA). Note that the CL shells are depleted in 18 O and 

13 

C relative to the NL shell (from Grossman et al., 

1993). 

Donets Basin brachiopods using the Ruhr method 

yielded an enormous δ 18 O range (-15 to -1‰), 

indicative of diagenetic alteration in meteoric waters 

(Bruckschen et al., 1999). The same samples 

prepared using the TAMU method had δ 18 O values 

equal to or higher than those sampled and 

analyzed by the Ruhr method, with an average 

difference of 3.1 ± 3.7‰ (N = 15; reported in 

Grossman et al., 2008, based on data in Bruckschen 

et al., 1999). These results imply that microsampling 

guided by PL and CL photomicrographs 

is more effective in avoiding diagenetic 

material than using milligram-sized samples of 

hand-picked shell crystals. A similar test with 

Russian Platform samples showed no significant 

difference between methods (Bruckschen et al., 

2001), suggesting that sample screening and sampling 

techniques are not critical for more pristine 

samples. Although more time consuming, thinsectioning 

and PL and CL microscopy has another 

49


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()*+, -../0*./01 

FIGURE 8.—Trace element concentrations of Carboniferous brachiopod shells and associated cement and matrix 

from North America (Mii et al., 1999). Filled and unfilled symbols differentiate averages of nonluminescent (NL) 

and luminescent (CL) spots for an individual specimen (error bars represent ±2 standard errors). MTE variability in 

NL shell includes the effects of seasonal environmental changes (Mii and Grossman, 1994). (A) S/Ca vs. Na/Ca. B) 

Mg/Ca vs. Sr/Ca. Note that CL shell is depleted in S, Na, and Sr relative to NL shell, whereas Mg content shows no 

systematic relationship with alteration as indicated by cathodoluminescence. Also note that trace element contents 

vary with genus (Mii et al., 1999). 

'"& 

advantage in that the samples are intact and available 

for studies with other proxies 

(photomicrographs can be viewed at 

http://geoweb.tamu.edu/faculty/grossman/SHELL 

_IMAGES/index.html). 

In general, the most diagenetically-resistant 

brachiopod material is the prismatic tertiary layer 

(e.g., Grossman, 1994; Lee and Wan, 2000). This 

material is less likely to show cathodoluminescence 

than fibrous secondary layer shell (Grossman 

et al., 1993). Unfortunately, use of the prismatic 

layer calcite complicates SEM screening for 

shell preservation because, unlike fibrous calcite, 

prismatic calcite has irregular microtextures that 

frequently resemble diagenetic cements (Bruckschen 

et al., 1999). Additional methods for the 

identification of diagenetic alteration in fossils are 

continually being sought. Two potential approaches 

are transmission electron microscopy 

(TEM), which can be used to study growth microfabrics 

(Ward and Reeder, 1993), and electron 

backscattered diffraction (EBSD), which tests for 

50


δ 18 O (‰ VPDB) 

δ 18 O (‰ VPDB) 

65 

75 

85 

95 

105 

115 

125 

Cretaceous 

Early Late 

Albian Cen Tu C S Camp Ma 

Aptian 

2 

1 

A 

0 

-1 

-2 

-3 

-4 

Shallow, Tropical- 

Subtropical 

-5 

-6 

3-7 

2-8 

B 

1 

0 

-1 

-2 

-3 

-4 

-5 

135 

H Ba 

Val 

Age (Ma) 

145 

155 

165 

175 

185 

195 

205 

215 

225 

235 

245 

Jurassic 

Triassic 

E Mid Late 

Early Mid Late 

I O An La Carn Norian Rhae H Sin Plies Toa Aa B B C Ox Kim Tith Ber 

COJa 

TJd 

10 20 30 40 50 60 

Shallow, 

Temperate 

10 20 30 40 

Isotopic temperature (°C) 

Isotopic temperature (°C) 

Trop/Sub. Temp. 

Brachiopod 

Belemnite 

Aragonitic mollusks (mean of serially-sampled shells) 

Shallow 

FIGURE 9.—Oxygen isotope records of tropical and temperate fossils and microfossils for the Mesozoic from 

Grossman (2012). Heavy lines represent running means with a 4 m.y. window and 2 m.y. steps, and fine lines show 

±1σ. COJa and TJd are the Callovian-Oxfordian (Jurassic) acme and Toarcian (Jurassic) decline respectively. Isotopic 

temperatures assume non-glaciated conditions (δw = -1‰ VSMOW). To correct for aragonite-calcite δ 18 O 

differences (Grossman and Ku, 1986), 0.6‰ is subtracted from the δ 18 O values of aragonitic taxa. Timescale from 

Gradstein et al. (2012). 

changes in crystallographic orientation (Pérez- 

Huerta et al., 2007). 

Isotopic records 

Mesozoic.—The δ 18 O compilations are shown 

in Figures 9, 10, and 11. Also shown are temperature 

scales using the Hays and Grossman (1991) 

reformulation of the O’Neil et al. (1969) equation 

51


assuming the non-glaciated reference state (δ 18 Οw 

= -1‰ VSMOW). Researchers have estimated the 

global ocean δ 18 O for an Earth without continental 

glaciers (ice-free or non-glaciated) using massbalance 

calculations based on the average mass 

and δ 18 O of the ocean and continental glaciers, 

which account for >99% of the water at the 

Earth’s surface. Early studies underestimated the 

magnitude of the effect of deglaciation because of 

inadequate information on the δ 18 O of Antarctic 

glaciers (see Pearson, this volume). With better 

data for modern glacial δ 18 Ow, estimates of the 

effect have converged on -1.0 to -1.1‰ (Shackleton 

and Kennett, 1975; Lhomme and Clarke, 

2005). Using Lhomme and Clarke’s (2005) estimate 

of -1.1‰ VSMOW for the glacial ice contribution 

and the average of LeGrande and 

Schmidt’s (2006) gridded data set (+0.01‰ 

VSMOW; Gavin Schmidt, pers. comm., 2012) for 

a global ocean δ 18 Ow yields -1.09‰ VSMOW for 

global seawater δ 18 O in a non-glaciated world. 

For simplicity, I use -1‰ VSMOW, the value used 

in many previous studies (e.g., Savin, 1977; Mii 

et al., 1999; Joachimski et al., 2004). The temperatures 

in Figures 9 through 11 and those presented 

in the text should be viewed with healthy 

skepticism as they do not consider geographic and 

long-term temporal variation in seawater δ 18 O 

(see earlier discussion of the influence of δ 18 O of 

environmental waters). Note that calculated isotopic 

temperatures for an interglacial ice-house 

world like our modern condition (δ 18 Οw = 0‰ 

VSMOW) will be ~5°C warmer than shown, and 

temperatures for a Pleistocene ice-age Earth 

(+1‰ VSMOW) will be ~10°C warmer. 

The Mesozoic data compiled in Figure 9 is 

from Grossman (2012), which is an update of data 

compiled in Veizer et al. (1999; 

http://mysite.science.uottawa.ca/jveizer/isotope_d 

ata/index.html) and Prokoph et al. (2008, Appendix 

A, Supplementary data). Much of the Triassic 

and Jurassic data come from Europe (England, 

Spain, France, Italy, and Poland; e.g., Anderson et 

al., 1994; Podlaha et al., 1998; Malchus and Steuber, 

2002; Jenkyns et al., 2002; Korte et al., 

2005b). The Triassic record is sparse, limited by 

availability of well-preserved marine fossils, and 

is based mostly on brachiopod shells. Oxygen 

isotope values for tropical brachiopods range 

from -6 to -0.5‰ and show an early Carnian 

(~225 Ma) increase of 2‰ that is attributed to 

cooling and to rising seawater δ 18 O due to increased 

evaporation (Korte et al., 2005b). Isotopic 

values for the latest Triassic based on tropical/ 

subtropical brachiopods average -1.5 ±1‰ (18 

±5°C), similar to δ 18 O values for early Jurassic 

belemnites from northern Europe (Jenkyns et al., 

2002). 

Northward movement of Europe during the 

Triassic and early Jurassic shifted samples from a 

tropical to temperate climate zone (Fig. 9). The 

mean δ 18 O values for temperate and tropical Jurassic 

belemnites are similar and increase in the 

early Jurassic to about -1‰ (~16°C) in the Pliensbachian, 

then sharply decrease in the Toarcian 

(TJd; ~181 Ma) to a minimum of about -3‰ 

(~30°C; Fig. 9). Middle Jurassic data are sparse 

but rise to a Callovian-Oxfordian acme (COJa; ca. 

165 Ma) of about 0.5‰ (~14°C), then decline to 

lower values of -1 to -1.5‰ in the late Jurassic 

(~17°C; 161–152 Ma). 

Oxygen isotope values of belemnites increase 

in the Early Cretaceous to a maximum of 0–1‰ 

(8–12°C) near the Valanginian-Hauterivian 

boundary (~136 Ma), interpreted as cooling (van 

de Schootbrugge et al., 2000; McArthur et al., 

2007). The δ 18 O values then decline to a minimum 

of -2 to -1‰ (16–20°C) in the middle Barremian 

(~128 Ma), interpreted as Barremian 

warming (Mutterlose et al., 2009). High belemnite 

δ 18 O values, sometimes equating to paleotemperatures 

less than 10°C in the early middle Jurassic, 

have been interpreted as reflecting a nektobenthic 

habitat (e.g., Dutton et al., 2007; Wierzbowski 

and Joachimski, 2007). This interpretation is supported 

by (1) low δ 18 O seasonality, (2) δ 18 O values 

similar to those of benthic foraminifera and 

bivalves, and (3) values lower than those of 

planktonic foraminifera and ammonites (Dutton et 

al., 2007; Wierzbowski and Joachimski, 2007, 

2009). 

Isotopic studies of fossils from the Cretaceous 

Western Interior Seaway of North America have 

yielded enigmatic patterns including anomalously 

low δ 18 O values in shallow-dwelling taxa such as 

nektonic mollusks and higher δ 18 O values in infaunal 

mollusks versus epifaunal ones (e.g., 

Wright, 1987; He et al., 2005). These patterns are 

believed to reflect a complex, salinity-stratified 

water column (e.g., Wright, 1987; Hudson and 

Anderson, 1989; He et al., 2005), or possibly 

submarine groundwater discharge (Cochran et al., 

2003). Since extracting global or regional climate 

information from these data is difficult, they are 

not shown in Figure 9. 

Paleozoic.—The first comprehensive Paleozoic 

δ 18 O record based on brachiopod shells was 

compiled by Ján Veizer and his colleagues and 

52


Age (Ma) 

250 

300 

350 

400 

450 

500 

550 

3 

Cambrian Ordovician Silurian Devonian Carboniferous Permian 

Gu Lo 

L P 

Late Mississippian Penn Cis 

Early M 

Fur Early Mid Late Llan W 

3 

2 

Te 

1 

δ 18 O brachiopod (‰ VPDB) 

-‐1 -‐3 -‐5 -‐7 -‐9 

Shallow, 

tropical/ 

subtropical 

brachiopods 

-‐11 26 

A B C 

ECi 

MLDd 

LSa 

LOa 

1 

10 20 30 40 50 60 70 

Isotopic temperature (°C) 

Shallow, 

temperate 

() and high 

latitude () 

brachiopods 

δ 18 O conodont (‰ VSMOW) 

24 22 20 18 16 

-‐1 -‐3 -‐5 -‐7 

δ 18 O (‰ VPDB) 

20 

30 40 50 

20 30 40 50 

Isotopic temperature (°C) 

14 

Shallow, 

tropical/ 

subtropical 

conodonts 

FIGURE 10.—Oxygen isotopic compositions of Paleozoic brachiopods (calcite) and conodonts (phosphate) from 

Grossman (2012). “Select” tropical-subtropical data are represented by open dark gray circles (see text for discussion); 

all other data are shown as open light gray circles. Unfilled boxes are brachiopod data for samples from accreted 

terranes in Japan and south China and represent open ocean conditions (box = range in values and bar = average; 

Brand et al., 2009). X symbols show clumped isotope temperatures from Came et al. (2007; Pennsylvanian and 

Silurian) and Finnegan et al. (2011; Silurian and Ordovician). Note that only Finnegan et al. data with Δ47 values 

above 0.589 are plotted (see Finnegan et al. for explanation). Phosphate data (C) are corrected to an NBS120c value 

of 22.6‰ (Vennemann et al., 2001; Joachimski et al., 2009; Pucéat et al., 2010). Only studies for which the value of 

NBS120c is given or determinable are used. Thick lines are running means with a 4 m.y. window and 2 m.y. steps. 

Light lines show values ± 1σ. Carbonate isotopic temperatures are based on the Hays and Grossman (1991) quadratic 

approximation of O’Neil et al. (1969) and phosphate isotopic temperatures are based on Pucéat et al. (2010), 

assuming seawater δ 18 O of -1‰ (VSMOW). LOa is the latest Ordovician acme, LSa is the late Silurian acme, 

MLDd is the mid-late Devonian decline, and ECi = early Carboniferous increase. Timescale from Gradstein et al. 

(2012). 

53


students (Veizer et al., 1997, 1999; Fig. 10). These 

data have been updated by Prokoph et al. (2008) 

and Grossman (2012). The prime features of the 

Veizer et al. (1999) curve are the steep decrease 

with age, roughly 1.3‰ per 100 m.y., and the 

high variability. Brachiopod δ 18 O values are generally 

10 to 4‰ for the Cambrian and Ordovician, 

8 to 2‰ for the Silurian and Devonian, and 7 to 

0‰ for the Carboniferous and Permian. Hypotheses 

to explain the decreasing δ 18 O with sample 

age are the same as those first proposed for δ 18 O 

trends in Precambrian rocks: (1) higher temperatures 

(e.g., Knauth and Epstein, 1976); and/or (2) 

lower seawater δ 18 O (Perry, 1967; Veizer and 

Hoefs, 1976) earlier in Earth history; or (3) the 

cumulative effects of meteoric diagenesis with 

age (Degens and Epstein, 1962). These hypotheses 

will be discussed later in this paper. 

Veizer et al. (1999) attribute the large scatter in 

the data to natural variability (~4‰ in tropical 

environments; Carpenter and Lohmann, 1995; 

Bruckschen et al., 1999) and inclusion of a small 

fraction of altered shells. The results of Bruckschen 

et al. (1999) discussed in Grossman et al. 

(2008), however, show that variability can be reduced 

by better sample screening and CL-based 

microsampling. With this in mind, Grossman 

(2012) reexamined and updated the compilations 

of Veizer et al. (1999) and Prokoph et al. (2008), 

culling samples that were not screened with 

cathodoluminescence microscopy or collected 

from strata with unusual fossil preservation. This 

approach reduces the variability considerably and 

allows better resolution of isotopic trends and 

events (Fig. 10A). 

The tropical/subtropical δ 18 O record for the 

Cambrian and Ordovician is overshadowed by 

questions of fossil preservation. This is especially 

true of Cambrian samples (Wadleigh and Veizer, 

1992). Isotopic studies of mid-Late Ordovician 

brachiopods (Qing and Veizer, 1994; Shields et 

al., 2003) show δ 18 O values increasing to a latest 

Ordovician acme (LOa; Hirnantian, ~445 Ma; 

Fig. 10A), when values increase from roughly 4‰ 

to between -2 and 0‰ before returning to preshift 

values (Marshall and Middleton, 1990; Qing 

and Veizer, 1994; Brenchley et al., 1994). This 

event of no more than a million years duration has 

been recognized in samples from Estonia, Sweden, 

North America and Argentina, and coincides 

with Hirnantian glaciation (Brenchley et al., 1995; 

Marshall et al., 1997). The isotopic shift (~3‰) 

equates to a temperature decline of 14°C. If one 

assumes ice volume changed from a non-glaciated 

state to an average late Pleistocene state (δ 18 Ow 

increase of +1.5‰), then the temperature decrease 

might only be 7°C (e.g., from 30°C to 23°C). 

Oxygen isotopic values decrease after the Hirnantian 

and into the Silurian. For most of the Silurian, 

δ 18 O values are relatively constant at -6 to -4‰, 

then increase to a late Silurian acme (LSa; Ludfordian, 

~420 Ma) of up to -2‰ (Fig. 10A). These 

results are based on brachiopods from Gotland, 

the Baltics, Scandinavia, Ukraine, Poland, and 

Anticosti Island, Canada (Samtleben et al., 1996; 

Wenzel and Joachimski, 1996; Bickert et al., 

1997; Azmy et al., 1998; Brand et al., 2006). The 

high late Silurian values do not correlate with any 

known glacial episode and, in combination with 

geologic data, have been interpreted as a decreased 

influence of freshwater input (Samtleben 

et al., 1996; Bickert et al., 1997). 

Many of the key features of the Ordovician 

and Silurian brachiopod δ 18 O records are mimicked 

in the δ 18 O records of conodonts. Ordovician 

brachiopods and conodonts both show increases 

from unusually low values to a maximum 

coincident with the Hirnantian glaciation 

(Brenchley et al., 1995; Marshall et al., 1997; 

Bassett et al., 2007; Trotter et al., 2008). Brachiopod 

and conodont paleotemperatures for the Late 

Ordovician yield warm to cooler temperatures of 

>32° to ~28°C (using Hays and Grossman, 1991, 

and Pucéat et al., 2010). Similarly, recent clumped 

isotope studies also register high temperatures for 

the Late Ordovician (32–37°C), cooling to 28– 

31°C in the Hirnantian (Fig. 10A; Finnegan et al., 

2011; also Affek, this volume). This convergence 

of brachiopod, conodont, and clumped-isotope 

temperatures argues for the verity of Late Ordovician 

isotope temperatures. 

Silurian brachiopod and conodont δ 18 O values 

(Fig. 10) show the same Llandovery increase, 

mid-Ludlow minimum, and late Ludlow acme 

(LSa; ~420 Ma; Wenzel et al., 2000). Conodont 

isotopic temperatures, first published as 24–33°C, 

similar to modern sea surface temperatures (assuming 

δ 18 Ow = -1‰), become 30–39°C with the 

Pucéat et al. (2010) 18 O paleothermometer, similar 

to brachiopod isotopic temperatures (24–41°C). 

Furthermore, recent clumped isotope studies of 

Silurian brachiopods argue for warm tropical 

temperatures (34–36°C) and seawater δ 18 O close 

to that of a modern non-glacial ocean (-1‰ 

VSMOW) (Came et al., 2007). Thus, clumped 

isotope studies support contentions of nearmodern 

seawater δ 18 O and retention of original 

oxygen isotopic compositions in well-preserved 

54


1 

0 

A 

10 

δ 18 O (‰) 

-1 

-2 

-3 

-4 

-5 

-6 

Arkansas 

IL, IN, IO, KS, OK, 

MS, NB 

New Mexico 

Texas 

Guadalupe Mtns., TX 

20 

30 

40 

Isotopic temperature (°C) 1 

Tournaisian 

Visean 

Mississippian 

Serp 

Bash Mos K 

Pennsylvanian 

Gz 

As 

Sakmar 

Artinsk 

Permian 

Kun R W Cap 

δ 18 O (‰) 

1 

0 

-1 

-2 

-3 

-4 

-5 

-6 

Fielding et al. (2008) 

B 

Glacial 1 Glacial II Glacial III 

Tournaisian 

Visean 

Mississippian 

C1 C2 C3 C4 P1 P2 P3 P4 

Serp 

Bash Mos K Gz 

Pennsylvanian 

Age (Ma, GTS2004) 

As 

Sakmar 

Artinsk 

Permian 

Isbell et al. (2003a) 

Kun R W Cap 

360 350 340 330 320 310 300 290 280 270 260 

10 

20 

30 

40 

Isotopic temperature (°C) 1 

1 

Assuming 

δw = -1‰ 

Moscow Basin (Bruckschen et al., 1999) 

Moscow Basin (Mii et al., 2001) 

Urals (Bruckschen et al., 2001; Korte et al., 2005a) 

Urals (Mii et al., 2001; Grossman et al., 2008) 

Urals (Popp et al., 1986) 

FIGURE 11.—Oxygen isotopic data for brachiopod shells from the North American Craton and the Russian Platform 

(modified from Grossman et al., 2008). Thick lines are running means for 3 m.y. window and 1 m.y. steps; 

thin lines represent ±1σ. Significant gaps in the record are shown as dashed lines. Data for the North American 

Craton are from Grossman et al. (1991, 1993), Mii et al. (1999), Korte et al. (2005a), and Grossman et al. (2008). 

Data for the Russian Platform are from Popp et al. (1986, as reported in Popp, 1986), Bruckschen et al. (1999, 

2001), Mii et al. (2001), Korte et al. (2005a), and Grossman et al. (2008). Isotopic temperatures assume nonglaciated 

conditions (δ 18 Ow = -1‰ VSMOW). Time-scale from Gradstein et al. (2004). 

brachiopod shells and conodonts. 

The tropical/subtropical brachiopod δ 18 O record 

for the Devonian is based on samples from 

USA, Spain, Germany, and China (Fig. 10A; 

Veizer et al., 1999; van Geldern et al., 2006), 

while the temperate record (~ latitude ≥ 35°) is 

based on samples from Morocco and Siberia, 

Russia (van Geldern et al., 2006). Tropicalsubtropical 

values rise to a Middle Devonian plateau 

of ~-3‰ (~25°C) then show a rapid mid-Late 

Devonian decline (MLDd) to a Givetian minimum 

of ~-6‰ (an uncomfortable ~40°C). The 

δ 18 O values remain mostly between -6‰ and -4‰ 

(40° and 30°C) during the Late Devonian. Though 

55


less detailed, the temperate record shows similar 

trends. These trends are interpreted in terms of 

temperature and seawater δ 18 O change, with cool 

temperatures in the Early and Middle Devonian 

and warm temperatures and lower seawater δ 18 O 

in the Late Devonian (van Geldern et al., 2006). 

The isotopic trends for Devonian conodonts (Fig. 

10C) are similar to those for brachiopods, with a 

mid-Devonian maximum and mid-Late Devonian 

decline (MLDd; Joachimski et al., 2004, 2009); 

but brachiopods show a 1–1.5‰ larger negative 

shift in the Givetian, resulting in high paleotemperatures 

of 30–40 °C (Joachimski et al., 2004). 

Using the Kolodny et al. (1983) 18 O paleothermometer 

for phosphate, Joachimski et al. (2004) 

obtained reasonable marine paleotemperatures of 

25° to 32°C for the Late Devonian, in contrast to 

30° to 40°C for brachiopod shells. However, the 

new 18 O paleotemperature relation of Pucéat et al. 

(2010) yields higher Late Devonian conodont paleotemperatures 

more in line with those for 

brachiopod shells (Fig. 10A). Thus, both materials 

give unusually high paleotemperatures for the 

Late (and Early) Devonian. These high values 

may represent a combination of temperature increase 

and moderate seawater δ 18 O decrease (van 

Geldern et al., 2006), perhaps reflecting circulation 

changes and greater influence of regional 

freshwater input. A lower global seawater δ 18 O of, 

for example, -2‰ VSMOW would yield an overall 

range of temperatures more palatable to biologists 

(20–35°C). 

The low brachiopod δ 18 O values of the Late 

Devonian continue into the earliest Carboniferous 

(Mississippian); values then increase through the 

Mississipppian (Fig. 10A; Popp et al., 1986; 

Veizer et al., 1986; Mii et al., 1999). Long Carboniferous 

records based on well-preserved shells 

are available from the North American Craton 

(NAC; especially the US mid-continent) and the 

Russian Platform (RP). The δ 18 O data from these 

regions are mostly between 0 and -5‰, yielding 

paleotemperatures mostly between 12° and 35°C 

for a non-glaciated Earth (seawater δ 18 O of -1‰ 

VSMOW), and 16° and 41°C for a moderately 

glaciated world (0‰ VSMOW). In contrast to the 

NAC and RP, many samples from central and 

western Europe have δ 18 O values lower than -6‰ 

(gray symbols, Fig. 10A; Bruckschen et al., 1999; 

Veizer et al., 1999) and yield variable isotopic 

temperatures often exceeding 50°C. These very 

low δ 18 O values, coeval with high values for other 

regions, undoubtedly reflect diagenetic alteration. 

Detailed examination of NAC brachiopod data 

shows: (1) a 3‰ rise in the Tournaisian to values 

of -2 to 0‰ (12–20°C); (2) a Visean decline to -4 

to -3‰ (25–30°C); and (3) a mid-Carboniferous 

increase of 1–2‰ with relatively constant Pennsylvanian 

values of -3 to -1‰ (16–25°C; Fig. 

11A; Mii et al., 1999; Grossman et al., 2008). Mii 

et al. (1999) attributed the mid-Carboniferous increase, 

also seen in the RP, to the initiation of 

continental glaciation (Fig. 11B; Bruckschen et 

al., 2001; Mii et al., 2001; Grossman et al., 2002, 

2008). 

The NAC and RP records show significant 

differences that call into question their global nature. 

One example, unusually low Uralian δ 18 O 

values for the late Serpukhovian (Fig. 11; Bruckschen 

et al. 1999), can be discounted because of 

exposure features suggesting diagenetic influence 

(P. Kabanov, pers. comm., 2007), but other NAC- 

RP differences appear to represent regional differences. 

These include the δ 18 O Moscovian- 

Kasimovian minimum and the Asselian maximum 

seen in RP, but not NAC data. The RP trends 

agree with the distribution of glacial sediments 

(Isbell et al., 2003; Fielding et al., 2008a) and 

may represent the record of global climate. This 

implies that local or regional variations in seawater 

δ 18 O influenced the NAC record. High North 

American δ 18 O values for late Tournaisian-early 

Visean brachiopods were originally interpreted as 

glaciation (Mii et al., 1999), but the distribution of 

NAC evaporites (Johnson, 1989) suggests that the 

18 

O enrichment was caused by regional aridification 

(Fig. 11A; Grossman et al., 2008). Another 

example of regional variation in seawater δ 18 O is 

the east–west δ 18 O increase from low values in 

Appalachian Basin (-3.8‰) to higher values in 

the Illinois Basin (2.4‰) and the US midcontinent 

(1.5‰; Flake, 2011). This trend likely 

represents increased influence of freshwater from 

the Appalachians, an interpretation supported by 

sedimentologic data (e.g., Cecil et al., 2003; Algeo 

and Heckel, 2008). 

The brachiopod and conodont δ 18 O records for 

the Carboniferous are similar in that both show an 

increase in the lower Mississippian (Fig. 10). 

However, brachiopod δ 18 O values decrease to 

about -3‰ in the late Visean, whereas conodont 

values continue to rise through the Mississippian 

to a late Mississippian maximum, before returning 

to more moderate values in the Pennsylvanian 

(Buggisch et al., 2008). Buggisch et al. (2008) 

observed positive δ 18 O shifts in the Tournaisian 

and Serpukhovian, and attributed them to major 

surges in glaciation. Slightly lower temperatures 

56


are indicated for Mississippian brachiopod shells 

(13–30°C) than for conodonts (17–32°C; Pucéat 

et al., 2010), except for the latest Mississippian, 

when conodont paleotemperatures are lowest. 

Permian brachiopods show an Asselian δ 18 O 

maximum in Uralian specimens, and a 

Sakmarian-Artinskian δ 18 O decline in Uralian and 

Australian specimens (Korte et al., 2005a, 2008) 

(Fig. 10A). These trends agree with sedimentologic 

evidence for an Asselian glacial acme and a 

Sakmarian-Artinskian decline (Isbell et al., 2003; 

Fielding et al., 2008a,b). Brachiopod δ 18 O values 

from the USA, Oman, and other Uralian sites fail 

to show the Asselian maximum and Sakmarian- 

Artinskian decline, instead showing increasing 

δ 18 O values in the Kungurian interpreted as reflecting 

aridification (Mazzullo et al., 2007; 

Grossman et al., 2008; Angiolini et al., 2009; 

Noret et al., 2009). High-latitude data from Australian 

brachiopods follow the pattern outlined in 

Korte et al. (2005a), but are offset by roughly 

+2‰, with values high in the Sakmarian (-0.2 

±1.2‰), decreasing in the early Artinskian (-1.8 

±0.7‰), then increasing in the Kungurian (-0.3 

±0.4‰) (Korte et al., 2008; Mii et al., 2012). Assuming 

the non-glaciated reference state (δ 18 Ow = 

-1‰ VSMOW), these values equate to a temperature 

trend from 12°C to 19°C to 13°C. Ivany and 

Runnegar (2010) suggested that Artinskian seawater 

δ 18 O around Australia might have been close 

to 4‰ based on an exquisite seasonal record from 

a serially sampled eurydesmid bivalve from Sydney 

Basin, Australia (see also Ivany, this volume). 

“Winter” values of 1‰ are interpreted to represent 

freezing temperatures based on the occurrence 

of glendonites and ice-rafted clasts, implying 

local seawater δ 18 O values of -3 or -4‰. 

Ivany and Runnegar (2010) believe that the influence 

of 18 O-depleted glacial meltwater was minimal, 

perhaps lowering δ 18 Ow by 1‰, and proposed 

a global mean δ 18 Ow closer to -2‰ than to 

the modern value of 0‰. This implies Artinskian 

tropical temperatures close to 18°C based on shallow 

brachiopod δ 18 O values (e.g., Korte et al., 

2005a; Grossman et al., 2008; Grossman, 2012). 

Alternatively, the low apparent seawater δ 18 O may 

be explained by low Artinskian ice volume (Fielding 

et al., 2008), higher eurydesmid growth temperatures 

(lack of winter growth), and/or greater 

meltwater influence. 

Isotopic studies of Permian conodonts further 

complicate the picture of late Paleozoic deglaciation. 

Based on 356 measurements of conodont 

elements from south China, USA, and Iran, Chen 

et al. (in press) observe relatively high and constant 

δ 18 O values (22.0–22.5‰) during much of 

the Cisuralian, equating to temperatures of 26°C 

to 30°C, assuming Pleistocene δ 18 Ow (+1‰; 

authors’ choice), and 22°C to 26°C, assuming 

modern δ 18 Ow (0‰). The δ 18 O values begin a 2‰ 

decline in the Kungurian, terminating in the early 

Wuchiapingian. Such a decline likely indicates a 

combination of warming and deglaciation, e.g., 4– 

5°C warming and 1‰ δ 18 Ow decrease equivalent 

to ~100 m sea-level rise. This contrasts with geological 

evidence suggesting that the main phase of 

late Paleozoic deglaciation occurred in the late 

Sakmarian (Isbell et al., 2003; Fielding et al., 

2008a,b). The conodont δ 18 O trend is well-defined 

for south China, but regional differences are observed 

such as higher Guadalupian values for 

Texas and lower Lopingian values for Oman. 

Lastly, conodont data for the latest Permian suggest 

a ~8°C warming trend in the latest Permian 

and across the Permian-Triassic boundary 

(Joachimski et al., 2012; Chen et al., in press). 

Long-term trends: Were early Paleozoic seas 

warm, 18 O-depleted, or is the trend a diagenetic 

artifact—Veizer et al. (1999) and more recently 

Jaffrés et al. (2007) and Prokoph et al. (2008) 

have highlighted the long-term increase of the 

sedimentary δ 18 O record throughout Earth history. 

Figure 10 also shows a temporal δ 18 O increase, 

though the δ 18 O trend for much of the late Paleozoic 

averages 1–2‰ higher than the Prokoph et 

al. (2008) curve. This, I believe, reflects better 

overall preservation of the samples used in the 

compilation. Supporting this contention is the 

convergence of brachiopod and conodont paleotemperatures 

discussed earlier. Conodont and 

brachiopod δ 18 O values in Figure 10 average 

about 2‰ lower for the late Ordovician to Devonian 

(18.3‰ and -4.6‰ respectively) compared 

with the Carboniferous and Permian (20.5‰ and 

-2.4‰). Assuming sample preservation can be 

ruled out as the cause of the 2‰ difference, then 

temperature must have decreased and/or seawater 

δ 18 O increased during the Paleozoic. 

In part, the higher Carboniferous–Permian 

δ 18 O values can be attributed to the transition 

from non-glaciated to icehouse conditions. Assuming 

a ~1‰ δ 18 Ow increase from ice-free to 

moderate icehouse conditions (Lhomme and 

Clarke, 2005), the remaining 1‰ difference could 

be explained by warmer early Paleozoic temperatures 

(~5°C), lower seawater δ 18 O values (to ~2‰ 

VSMOW), or some combination of the two. Assuming 

the non-glaciated reference state (1‰ 

57


VSMOW) yields mean isotopic temperatures for 

Late Ordovician through Devonian brachiopods 

and conodonts as high as 41°C (Fig. 10). Biologists 

and paleobiologists generally consider the 

thermal limit of metazoan life to be roughly 38°C 

(Brock, 1985). The thermal limit for aquatic mollusks 

is in the 33–38°C range (Hicks and McMahon, 

2002), while some worms from hydrothermal 

vents have thermal limits above 45°C (Pörtner, 

2002; Lee, 2003). Early Paleozoic brachiopods 

may have been better adapted for high temperatures 

than modern species, but this is difficult 

to prove. As discussed earlier, clumped isotope 

results suggest temperatures of 32–37°C in the 

Late Ordovician (Finnegan et al., 2011) and 34– 

36°C for the Silurian (Came et al., 2007), and little 

change in seawater δ 18 O (Fig. 10A); however, 

the susceptibility of clumped isotope signatures to 

reordering and the applicability of the Ghosh et al. 

(2006) paleothermometer to brachiopod shells are 

still open questions (e.g., Passey et al., 2011; 

Henkes et al., 2012). 

Numerous studies have examined the evolution 

of seawater δ 18 O through time using measurements 

of various materials and mass-balance 

models (see Muehlenbachs, 1998, and Jaffrés et 

al., 2007, and references therein). In general, 

ophiolites, ore deposits, meteoric cements, and 

fluid inclusions show no evidence of a temporal 

increase in δ 18 O (Muehlenbachs, 1986; Gregory, 

1991; Hays and Grossman, 1991; Knauth and 

Roberts, 1991; Muehlenbachs, 1998), but the data 

are too variable or temporally limited to detect 1– 

2‰ changes. The model results fall into two categories: 

those showing that seawater δ 18 O is buffered 

by crustal processes with variance within ±1- 

2‰ (e.g., Muehlenbachs and Clayton, 1976; 

Gregory, 1991; Lécuyer and Allemand, 1999), and 

those arguing for increases as large as 5‰ since 

the Late Ordovician (e.g., Wallmann, 1991; Jaffrés 

et al., 2007). While a critical review of these 

models, their parameters, and their assumptions is 

beyond the scope of this chapter, the data presented 

here suggest that the seawater δ 18 O increase 

in response to crustal cycling was


global climate change. 

Clumped isotope paleothermometry has the 

potential to solve the half-century debate over 

whether the low δ 18 O of Precambrian rocks and 

early Paleozoic fossils represents warm temperatures, 

changing seawater δ 18 O, or sample alteration. 

Early results with clumped isotopes suggest 

warmer temperatures in the early Paleozoic, 

though slightly lower seawater δ 18 O (e.g., 2‰ 

VSMOW) cannot be ruled out. If these conclusions 

endure, paleobiologists will have to rethink 

temperature tolerances of metazoans and the role 

of climate change in evolution. 

Lastly, researchers are gaining a renewed appreciation 

for the special character of the epicontinental 

seas on which the Paleozoic and early 

Mesozoic isotopic and biologic records are based. 

Armed with an arsenal of geochemical (e.g., trace 

elements and C and Nd isotopes) and geologic 

approaches, researchers are developing a better 

understanding of when these seas reflect openocean 

conditions, and when they do not. These 

studies are enabling a more accurate picture of 

global climate, and improving our understanding 

of the control of the physical environment on biogeography 

and evolution. 

ACKNOWLEDGMENTS 

I thank the US National Science Foundation 

for previous and current (EAR-0643309) support 

of our isotopic studies of late Paleozoic climate 

and oceanography. The manuscript was improved 

by the helpful reviews of Linda Ivany and Jasmine 

Jaffrés, careful proofreading by Lauren Graniero, 

and helpful discussion with Bryan Bemis, 

Sang-Tae Kim, Gavin Schmidt, and Howie Spero. 

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