ELSEVIER Chemical Geology 125 (1995) 73-99

lMuuDlNG

ISOTOPE GEOSCIENCE

Sr and Nd isotopes at the Permian/Triassic boundary: A record of climate change

E.E. Martin 2y* , J.D. Macdougall Scripps Insrirubon of Oceanography, University of California, San Diego, La Jolla, CA 92093-0220, USA

Received 8 July 1994; accepted 1 March 1995 after revision

Abstract

We present a detailed curve of seawater 87Sr/86Sr for the Middle Permian to Triassic based on analyses of conodonts from overlapping sections in the U.S.A. and Pakistan, correlated using conodont biostratigraphy. The isotope ratio decreased in the Middle Permian at an average rate of 0.000062 Ma- I, reached a minimum in the Capitanian (257-258 Ma), and increased in the Late Permian at an average rate of 0.000097 Ma-‘. The Late Permian rate of increase was roughly two and a half times greater than the average increase over the past 40 Ma, and approximately equal to the highest Cenozoic rates, which occurred over much shorter time intervals.

Modeling results suggest that decreasing Middle Permian 87Sr/86Sr ratios were driven by changes in the riverine Sr flux to the oceans, while increalsing ratios in the Late Permian/Triassic are attributed to both increased riverine “Sr/‘%r and flux. The reduced Middle Permian riverine flux coincides with extreme continental aridity associated with the formation of Pangea and recorded by massive evaporite deposits. In addition, mountains in the equatorial region of Pangea may have created a rain shadow, thereby minimizing precipitation in regions that currently contribute the bulk of chemical weathering products to the ocean.

Increasing riverine 87Sr/86Sr in the Late Permian is suggested by the observation that ‘43Nd/ 144Nd values decrease at the same time; however, the source of radiogenic Sr is not known. Frequently cited mechanisms for increasing 87Sr/86Sr in runoff, such as glaciations and continent-to-continent collisions, coincide instead with decreasing seawater 87Sr/86Sr in the Middle Permian. One possible source may have been deep erosion into older orogens, associated with a dramatic increase in chemical weathering in the Late Permian. The cause of enhanced weathering appears to have been increased levels of atmospheric CO2 with associated global warming and increased humidity. Proposed sources of COZ include dissociation of gas hydrates and oxidation of organic matter during extreme sea level regression, as well as volcanic emissions from Siberian Traps eruptions. Continental floral and fauna1 distributions are consistent with this interpretation, as are oceanic 613C patterns and variations in shallow-water sediment lithologies.

1. Introduction

The period of time near the Permian/Triassic (P/ Tr) boundary was marked by a number of anomalous

[=I * Corresponding authcr. * Present address: 1112 Turlington Hall, Department of Geology,

University of Florida, Gainesville, FL 3261 l-7340, USA.

events in the history of the Earth. The magnitude of many of these events, including the biotic crisis, sea level regression and seawater geochemical anomalies, is unsurpassed during Phanerozoic time. The geochem- ical anomalies include dramatic variations in seawater S34S (Claypool et al., 1980), 613C (Baud et al., 1989) and *‘Sr/*%r (Burke et al., 1982; Denison et al., 1994))

0009-2541/95/$09.50 (0 1995 Elsevier Science B.V. All rights reserved SSDIOOO9-2541(95:100081-X

74 E.E. Martin, J.D. Macdougall /Chemical Geology 125 (1995) 73-99

suggesting major chemical changes in the oceans. Spe- cifically, the seawater 87Sr/86Sr curve for this interval reaches a minimum value for the Phanerozoic (Burke et al., 1982; Popp et al., 1986b; Holser and Magaritz, 1987; Denison et al., 1994). Changes before and after this minimum appear to be rapid, comparable to the well-documented rate of increase over the past 40 Ma.

variation, nor the cause(s) of the extinction event. These may or may not be related.

This paper evaluates the timing, magnitude and rate of change of seawater *‘Sr/‘%r for the time interval spanning the P/Tr boundary. Temporal variations in this isotopic ratio reflect changes in the riverine (con- tinental) and hydrothermal inputs to the ocean. Specif- ically, the isotopic signature and magnitude of the continental flux are controlled by the interaction between tectonics and climate. Tectonic uplift and landmass distribution determine the age, type and quan- tity of rock exposed to weathering, while climate con- trols the intensity and rate of weathering. In contrast, the hydrothermal flux is presumably controlled by the rate of sea-floor spreading.

We evaluate the P/Tr Sr data from both a forward and inverse perspective. On the one hand, we use the timing, magnitude and rate of change of seawater s’Sr/ ‘?Sr variations to constrain the processes occurring dur- ing the Permian. On the other, we apply our knowledge of geologic and environmental conditions at that time to evaluate the Sr cycle.

Many models of the seawater Sr isotope system have focused on the rapid increase in s7Srla6Sr during the Cenozoic (Palmer and Elderfield, 1985; Hodell et al., 1989; Capo and DePaolo, 1990; Hodell et al., 1990; Richter et al., 1992). The general increase has been attributed to tectonics (Raymo et al., 1988; Edmond, 1992; Raymo and Ruddiman, 1992; Richter et al., 1992), while climate has been proposed as the cause of some of the shorter time-scale changes within the general increase (Hodell et al., 1989, 1990; Capo and DePaolo, 1990; Miller et al., 1991; Zachos, 1993). Recently, Edmond ( 1992) proposed that Himalayan- style collisional tectonics are required to produce dra- matic large-scale increases in 87Sr/86Sr because this is the only mechanism whereby radiogenic Sr can be redistributed into easily weathered phases.

Much of the data in the literature for paleoseawater 87Sr/86Sr come from analyses of carbonate sediments sampled at widely distributed localities. In contrast, all of the data presented in this paper are from conodonts collected from a series of overlapping sections corre- lated by their conodont biostratigraphy. Conodonts are teeth-like elements from an extinct soft-bodied chor- date group (Briggs et al., 1983; Aldridge, 1987) and are composed of francolite, a carbonate fluorapatite, which is believed to be less susceptible to diagenetic alteration than carbonate (Bemat, 1974; Kovach, 1980; Staudigel et al., 1985; Bertram et al., 1992; Ruppel et al., 1993). Another advantage is that carbonate fluor- apatite incorporates high concentrations of both Sr and Nd (Wright et al., 1984; Shaw and Wasserburg, 1985). Conodonts evolved rapidly, therefore their biostrati- graphic zones cover short time intervals and relative sample ages are well defined.

In addition to the Sr data, we present detailed records of seawater ‘43Nd/‘44Nd evolution for the Proto-Pacific (Panthalassa) and Tethys oceans in the Permian. These data impose added constraints on the sources of sea- water a7Srla6Sr variations.

2. Conditions and events during the Permian

The rapid rise of seawater s7Srla6Sr in the Late Per- Information about the tectonic and climatic condi- mian to Triassic is different from the increase over the tions for the Permian through Triassic creates the past 40 Ma in several respects. First, the Permian is one framework for the evaluation of Sr isotopic variations of the least radiogenic intervals for seawater Sr during in seawater. The timing of many of the major events of the Phanerozoic. Second, it is a period during which this time period is summarized in Fig. 1. As already the model of Richter et al. ( 1992), which accurately mentioned, the Permian was a unique period in Earth’s reproduces many sections of the Phanerozoic 87Sr/86Sr history in many respects. The final consolidation of curve, does not match the measured data. Finally, it is Pangea created a supercontinent that probably experi- a major extinction boundary, likely to be characterized enced severe climatic conditions (Crowley et al., 1989; by dramatic environmental changes that could affect Crowley and North, 1991; Crowley, 1994), further the Sr cycle. We would like to emphasize that the focus intensified by extremely low sea level stands. Global of this paper is to identify the cause(s) of the isotopic sea level curves by Hallam ( 1984), Holser and Magar-

E.E. Martin, J.D. Macdougall /Chemical Geology 12.5 (1995) 73-99

Bee

.u 247.1

3 .” $

-251 .O

stecle

Scythian

E Changsingian

*g 253.5 c s*

8 .E P 3 Wuchiapingian

g! 256.0 I- 3 Capitanian 87Sr/86Sr minimum

S -2156.5 I

Wordian I N. Kazanianl E I a,

Kungurianl Leonardian I

s ._ E ; 0 )I Z m 287.5 w

--496.0

Carboniferous

I

- Rq -

r.

Fig. 1. Timing of climatic, tectonic and geochemical events during the Permian. Data from: ’ Frakes ( 1979), and Caputo and Crowell ( 1985); ’ Epshteyn (1981); 3 Ziegler ( 1989), and Scotese and McKerrow (1990); 4 Nalivkin (1973), Zonenshain et al. (1984), and Scotese and McKerrow (1990); 5 Baksi and Farrar (1991), Dalrymple et al. (1991). and Renne and Basu (1991); 6 Zharkov (1981); ’ Hallam (1984), Holser and Magaritz ( 1987), Ross and ROSS ( 1987)) and Wignall and Hallam (1993); 8 Popp et al. (1986a), Magaritz et al, (1988). Baud et al. ( 1989), Gruszczynski et al. ( 1989). and Holser et al. (1989); ’ Claypool et al. (1980); lo Delaney et al. (1989); ‘I Sepkoski (1982) cited by Holser and Magaritz (1987), and Raup and Sepkoski (1986); ‘* Erwin ( 1993). The time scale used in many of these publications varied slightly from the one presented here. In these cases, we made our best effort to correlate between the data sets. As a result, the age ranges presented here do not always match the ranges in the original publications.

itz ( 1987), Ross and Ross ( 1987), and Hallam ( 1992) outline a regressive phase from the Early Permian to the P/Tr boundary, with an abrupt drop at, the bound- ary, or just before the boundary with transgressive con- ditions at the boundary (Wignall and Hallam, 1993) (Fig. 1). According to Hallam ( 1984) this lowstand represented the lowest sea level during the entire Pha- nerozoic. Estimates of the magnitude of the regression range up to 210 m (Forney, 1975) or even 280 m (Holser and Magaritz, 1987). In contrast, the sea level

curve by Vail et al. ( 1977) indicates general regressive conditions from Mid-Carboniferous to Jurassic, with the lowest point occurring about Mid-Permian. This interpretation is less consistent with geologic evidence of a worldwide unconformity at or near the P/Tr boundary.

Gradual sea level variations are commonly attributed to changes in the volume of ocean ridges (Hays and Pitman, 1973). In the Permian this could have occurred by slower spreading rates or actual loss of ridge seg-

76 E.E. Martin. J.D. Macdougalll Chemical Geology 125 (1995) 73-99

ments. The assemblage of a single huge continent implies that many of the plate boundaries were conti- nent-to-continent. As Schopf (1974) has suggested, this might reduce the free movement of plates and con- sequently decrease spreading rates. Consumption of ocean basins during continent-to-continent collisions would also destroy segments of the mid-ocean ridge (Hallam, 1977).

The Early Permian marked the end of the Late Pale- ozoic ice age, which began in the Early Carboniferous when a portion of Gondwana was located at the south pole. During this glaciation extensive ice sheets cov- ered South America, south-central Africa, India, Ant- arctica and Australia (Caputo and Crowell, 1985). Gondwana migrated north during the Late Carbonif- erous and Early Permian, and most of the ice had melted by Sakmarian time (Frakes, 1979) (Fig. 1). Final dropstones are recorded from Australia and Antarctica in the lowermost Kazanian (throughout this paper we refer to the stages by their Russian names only for the sake of brevity) (Caputo and Crowell, 1985). Minor Kazanian-aged glacial dropstones have also been iden- tified in the northern hemisphere on the Kolyma block of Siberia (Epshteyn, 198 1). This cooling reflects the encroachment of Pangea on the north pole. Climate models predict conditions suitable for extensive glacia- tion at both poles during the latter part of the Permian (Crowley et al., 1989; Kutzbach and Gallimore, 1989). The absence of substantial ice development led Erwin (1993) to speculate that atmospheric CO, concentra- tions may have been high at that time, resulting in global warming.

For much of the Permian extreme continentality gov- erned the climate in the interior of the supercontinent. Climate models predict daytime summer temperatures as high as 35~lS”C (Crowley et al., 1989; Crowley, 1994) with seasonal variability of 30°C; a range cur- rently experienced only in Siberia and northern Canada (Crowley and North, 1991), although the absolute tem- peratures in these areas are lower. Arid conditions pre- vailed in the interior regions, with estimates of precipitation at N 50% of the modem value (Kutzbach and Gallimore, 1989; Crowley, 1994). In contrast, strong monsoonal circulation may have delivered sea- sonal rains along the eastern continental margins (Rob- inson, 1973; Kutzbach and Gallimore, 1989). Parrish (1993) suggests that the intensity of monsoonal cir- culation increased from the Late Carboniferous to a

maximum in the Triassic. As Robinson ( 1973) points out, the distribution of climate-sensitive rocks, such as coals, eolian deposits, laterites, red beds and evaporites is consistent with these climate model results.

Massive evaporite beds were deposited in the Middle and Late Permian and the Mid-Triassic (Gordon, 1975). In a detailed analysis, Zharkov (1981) esti- mated the volume of salts and sulfate rocks preserved from each stage of the Paleozoic (Fig. 1). He con- cluded that one third of all the salt and sulfate rocks deposited in the Paleozoic were deposited during the interval he labels Kungurian. Another 6% of the Pale- ozoic salts and sulfate rocks are preserved in the Middle to Late Permian stages of the Kazanian and Tatarian. Fischer ( 1964) and Stevens ( 1977) suggested that the Middle to Late Permian ocean may have been stratified because of this extensive evaporite deposition. They speculate that dense brines, formed during evaporite precipitation, could have been stored in the deep ocean, while surface water salinities were reduced by as much as 3-3.5%0. Both authors discussed this phenomenon in terms of the effects of brackish conditions on shal- low-water fauna1 extinctions. Holser and Magaritz (1987) noted that the associated reduction of the Sr concentration in shallow-water habitats could account for the very rapid changes in 87Sr/86Sr observed in Late Permian samples. However, fluid inclusion data from Horita et al. ( 1991) indicate that salinities in shallow- water environments were similar to modem seawater values.

Zharkov ( 1981) calculated that 1.48. lo5 km3 of sulfate rocks are preserved from the entire Permian. At 1000 ppm Sr, 3.62. lOi mol Sr would be stored in these rocks, equivalent to 3% of the modern ocean’s Sr reservoir. Evaporites are highly susceptible to weath- ering, thus the original volume deposited may have been much greater. Stevens (1977) theorized that the weathered fraction could easily be 50%, implying that the Permian evaporites may have removed N 6% of the Sr from the ocean. This is a relatively minor change and is consistent with the fluid inclusion data (Horita et al., 1991).

Isotope curves for Si3C and S34S also record large variations in the Permian. In general, Late Carbonif- erous to Early Permian seawater seems to have been characterized by high values of S13C, ranging up to u + 4 to + 6%0 (Poppet al., 1986a; Holser and Magar- itz, 1987), with intermittent periods of lower Si3C-

E. E. Martin, J.D. Macdougall / Chemical Geology 125 (1995) 73-99 II

values. A final high spike preserved globally in the Tatarian (Holser et al., 1986; Baud et al., 1989) con- trasts with generally decreasing values from the Kaza- nian to the P/Tr boundary, with a particularly rapid decrease at the boundary (Fig. 1) (Baud et al., 1989; Gruszczynski et al., 1989; Holser et al., 1989). Con- stant A ‘3C-values ( 6’3C,,b - SL3C,) spanning the P/ Tr boundary suggest that the carbon isotopic variations record a global, whale ocean signal (Magaritz et al., 1992). Erwin ( 1993‘) describes three possible sources of light carbon which are consistent with conditions at that time: ( 1) an influx of juvenile carbon from vol- canic degassing (8°C = -5%0) during the Siberian Traps eruptions, (2) exposure, weathering and oxida- tion of buried organic carbon ( SL3C = - 25 to - 20%0) during the Late Permian sea level regression, or (3) release of methane ( 613C = -65%0) during gas hydrate dissociation related again to the sea level regression. Mass-balance calculations indicate that the volcanic input would only have a minor effect, but either of the other two sources would be capable of producing the observed decrease (Erwin, 1993). In addition, the relatively high 613C-values preceding the decrease could be related to enhanced burial of organic carbon relative to carbonate.

Sulfur isotopic variations recorded in sulfates are even more dramatic (Claypool et al., 1980). S34S decreases gradually throughout the entire Permian (Fig. 1) from a value of - + 14%0 at the Permo- Carboniferous boundary to the Phanerozoic minimum of - + 11%0 in the earliest Triassic, then increases rapidly to - +26 or + 28%0 in the Scythian. The decreasing values probably represent a period of net pyrite oxidation and. erosion, while the rapid increase indicates sulfide removal through pyrite burial. To address the very rapid increase in the Early Triassic, Holser (1977) proposed catastrophic mixing between the surface ocean and a peripheral basin containing elevated 634S-values created by pyrite deposition.

Tectonic events may have a significant influence on seawater 87Sr/86Sr. For example, evidence suggests that the rapid rise over the last 40 Ma is related to uplift and unroofing of the Himalaya (Raymo et al., 1988; Hodell et al., 1989, 1990; Palmer and Edmond, 1989; Edmond, 1992; Krishnaswami et al., 1992; Raymo and Ruddiman, 1992; Richter et al., 1992). The two major continent-to-continent collisions around Permian time are reflected in the Hercynian and Uralian Orogenies

(Fig. 1) . Clockwise rotation of Gondwana led to a NE to SW progression of continental collisions along the Hercynian megasuture (Scotese and McKerrow, 1990) which peaked in the Late Carboniferous (Ziegler, 1989) and ended in the Early Permian. The Uralian collision between Baltica and Kazakstan also began in the Carboniferous (Nalivkin, 1973; Zonenshain et al., 1984). In this case, the main collision event occurred in the Early Permian (Nalivkin, 1973; Scotese and McKerrow, 1990). Sediment analysis indicates that the Urals were high, snow-capped peaks in the Sakmarian to Artinskian. Thin fanglomerate and sandstone sequences suggest that compression had essentially ceased by Kungurian time. And by the Late Permian the Urals are described as “low and eroded” (Naliv- kin, 1973). Thus, most of the effect of these orogenies on seawater 87Sr/86Sr should have occurred during the Late Carboniferous and Early Permian.

Massive flood basalt volcanism produced the Sibe- rian Traps around P/Tr time (Fig. 1). This may rep- resent the largest flood basalt province in the Phanerozoic, with an estimated original volume of > 1.5 ?? lo6 km3. The eruption interval for this event appears to have been very short, possibly < 1 Ma (Renne and Basu, 1991). Recent attempts to date the Siberian Traps with high-precision techniques yielded ages roughly equivalent to the P/Tr boundary (Baksi and Farrar, 1991; Dahymple et al., 1991; Renne and Basu, 1991). Campbell et al. ( 1992) found that zircon 206Pb/238U ages for rocks comagmatic with the Sibe- rian Traps and for the boundary clay in a Chinese sec- tion agreed within analytical uncertainty.

Analysis of the boundary clay in China indicates that it may be altered ash from a silicic volcanic event that was unrelated to the Siberian Traps (Zhou and Kyte, 1988). The timing of both the large-scale basaltic and smaller-scale silicic volcanic eruptions coincided approximately with the end of a period of stable reversed magnetic polarity known as the Kiamen Long Reversed Superchron (Fig. 1). The relationship between field reversals, massive volcanism and extinc- tion events is unclear; however, it appears that the peak of the gradual extinction event in the Late Permian and most of the volcanism occurred several millions of years after the reversal (Fig. 1) (Raup and Sepkoski, 1986; Holser and Magaritz, 1987).

78 E.E. Martin, J.D. Macdougall/Chemical Geology 125 (1995) 73-99

3. Stratigraphic background

The rate of isotopic variation in seawater that is cal- culated from the measured data is highly dependent on the time scale applied. Unfortunately, there is little con- sensus on the absolute ages, or even the nomenclature, for Permian stages. This difficulty can be attributed to the extremely low sea level and low fauna1 diversity and abundance at that time. Particularly in the Late Permian, sediments were frequently deposited in iso- lated basins and the high proportion of endemic fauna complicate inter-basin correlations.

We adopted the sequences of stages outlined in War- dlaw ( 1995). Assigning absolute ages to biostrati- graphically defined stages is a rather uncertain process. The time scale adopted here is modified from Harland et al. ( 1990). Since the completion of that time scale, ion microprobe dates for zircons from a bentonite boundary layer at the Meishan section, Changsing, China, have been published, which give an age of 251.2 f 3.4 Ma for the P/Tr boundary (Claod-Long et al., 1991) . This is significantly older than the Harland et al. ( 1990) estimate of 245 Ma. The number of radi- ometric ages from the Permian is limited, and the impli- cations of the older period-boundary date for the timing of the stage boundaries is unclear. In general we have maintained the length of each stage as given in Harland et al. ( 1990), but shifted their boundaries by 6 Ma; however, our Kazanian and Ufimian stages are 1 Ma longer and the Kungurian and Artinskian stages are 1 Mashorter (B.R. Wardlaw, pers. commun., 1993) than those of Harland et al. ( 1990). With the exception of the Kazanian/Ufimian boundary, our age assignments fall within the minimum and maximum values plotted on the chronograms in Harland et al. (1990). Sweet ( 1992) suggested that the Late Permian may have been much shorter than is assumed in this paper. In this case, rates of isotopic change over this interval would be even more rapid than we have calculated.

No single stratigraphic section provides a complete, unaltered, Permian through Triassic marine sequence with abundant conodonts. Therefore, we pieced together several overlapping sections from the U.S.A. and Pakistan for this study (Fig. 2). Sample density is greatest for the Middle and Late Permian, with just a few samples representing older Permian marine sedi- ments. All of the conodonts for this study were very generously provided by B.R. Wardlaw (U.S.G.S., Res-

ton, Virginia) who collected, processed and identified the specimens.

The completeness of various P/Tr sections is still disputed (for details see Sweet et al., 1992), but there is general agreement that the most complete are from the Tethyan realm. Uppermost Permian and Lower- most Triassic samples for this study are from the Salt and Kbisor Ranges in Pakistan (Wardlaw and Pogue, 1995). These specimens were collected from the Amb, Wargal, Chhidru and Mianwali formations, which range in age from Kazanian to Smithian (Fig. 2). As pointed out by Teichert ( 1990) and Sweet ( 1992), the Uppermost Permian is missing from these locations. However, data from above and below the missing inter- val suggest that 87Sr/86Sr increasescontinuously across the boundary. Detailed descriptions of the formations are found in the report of the PJRG ( 1985) and War- dlaw and Pogue ( 1995). In general, the sediments are composed of interbedded limestone, dolomite, marl, sandstone and shale, representing deposition in inter- tidal, marginal marine and shallow marine settings. All of the conodonts were collected from limestone inter- vals which are believed to have been deposited under normal marine conditions.

Pakistani sections overlap with U.S.A. sections in the Kazanian, Capitanian, Smithian and Dienerian. Most of the U.S. material is from the Glass and Del Norte Mountains in southwest Texas (Wardlaw and Grant, 1990; Wardlaw et al., 1990, 1991). The samples range in age from Artinskian to Capitanian (Fig. 2), and were collected from the Skinner Ranch, Cathedral Mountain, Road Canyon, Word, Vidrio and Altuda for- mations. These sediments were deposited at the south- ern margin of the Permian Basin in shallow intertidal or lagoonal to deeper-water shelf edge and slope envi- ronments.

Additional U.S.A. samples come from the Baldwin Creek section at the eastern edge of the Phosphoria Basin in Wyoming, which includes the Ervay, Franson and Grandeur members of the Park City Formation and the Retort Member of the Phosphoria. It has been pro- posed that all of these sediments were deposited in a carbonate ramp setting along the Cordilleran margin (Wardlaw and Collinson, 1986). Three Early Permian samples are from the Garden Valley Formation in cen- tral Nevada, which was deposited along the Cordilleran margin in shallow shelf to outer shelf settings (Galle- gos and Wardlaw, 1992). These specimens are dated

E. E. Martin, J.D. Macdougall / Chemical Geology 125 (1995) 73-99 19

United States Pakistan

AtIe

Q 247.1 247.9 Spathian

a Smithian

g 249.4 Dienerian

Greisbachian -,251.0

: Changsingian

E 253.5

t

5 ‘Z

0 m ;i

Wuchiapingian

g 256.0 k 4 Capitanian

--256.5

g Kazanianl

E Wordian

2 262.0 sl =

Ufimianl Roadian

4 264.1 z Kungurianl

Leonardian --2%

nlne !scaIO change

Arlinskian

274.6

3 E &

Sakmarian

P * 5 W 207.5

--296.0

Carboniferous

1 5

;N

-

-

-

-

-

-

-

-

-

-

t

m B B

g -

L k

-

-

-

-

-

-

-

-

Fig. 2. Proposed Permian time scale and age distribution of conodont samples from the U.S.A. and Pakistan. The time scale is modified from Harland et al. ( 1990). B.R. Wardlaw defined the conodont biostratigraphy.

as Late Asselian, Early Sakmarian and Early Artin- skian. Because there are long, unsampled time intervals between each specimen, the relative ages are not well constrained. The lowermost Permian is represented by two samples from just above the Permo-Carboniferous boundary in the Mid-Continent Basin, Kansas. They are from a basal marine regressive limestone in the classic cyclothem sequences of the Neva Limestone. Triassic material from the U.S.A. comes from the Crit-

tenden Springs section in northeastern Nevada. These sediments from the Thaynes and Dinwoody formations represent deposition in a lagoon to high-energy near- shore setting along the Cordilleran margin (Carr and Paull, 1983).

Absolute errors for radiometrically dated material from the Permian range from It 2 to f 10 Ma (Forster and Warrington, 1985; Harland et al., 1990). This rep- resents a minimum uncertainty because the location of

80 E.E. Martin, J.D. Macdougall /Chemical Geology I25 (1995) 73-99

(4

+ -

Fig. 3a i

250 260 270 280 290 300

AgO (MaI (b) 0.7065 , , , r .

?? h

3i % 3

0.7075 -

m

0.7070 -

Fig. 3b

0.7065 . ’ ’ ‘,’ ’ ’ ,’ I Triassic

Late pennian Middle Permian Earlv Ch 1 Wu 1 Ca 1 Ka 1 Uf 1 Ku 1 Ar

245 250 255 260

Age (Ma)

265 270

(c) 0.7065 . . I . . I I 1 . I . m @.,

. e@ : ;;

0.7080 - \,j :.e:’ -

6

2

p 0.7075 -

OD

0.7070 -

0.7065 /I\ 245 250 255 260

Age VW

265 270

Fig. 3. Permian/Triassic seawater s7Sr/*% vs. age. a. A comparison of Permian/Triassic *‘Sr/% data vs. age from this study (0 = U.S.A. and ??= Pakistan), as well as A = Veizer and Compston ( 1974); 0 = Popp et al. ( 1986b); H = Nishioka et al. (1991); and + = Denison et al. (1994) for the Permian data and Koepnick et al. ( 1990) for the Triassic data. All ratios were corrected for interlaboratory bias to an NBS 987 value of 0.710269. Error bars are smaller than the size of the symbols.

many of the dated samples relative to the stage bound- aries is imprecisely known (for further discussion see Odin, 1985). For correlation between sections the pri- mary concern is relative age estimates. The samples can be placed quite confidently within biostratigraphic conodont zones; thus an accurate statement of the rel- ative error in age would be the average length of a conodont zone, which is remarkably short due to the rapid evolution of the conodont organism. For material younger than Sakmarian this value is N 2 Ma, which is consistent with the spread of data points along the age axis on the seawater Sr isotope curve (Fig. 3). For Sakmarian and Asselian material y 5 Ma appears to be a more accurate estimate.

Thermal maturation of organic material incorporated in conodont apatite results in progressive color altera- tion of the specimen. Color Alteration Index (CAI) values of 1 indicate that the conodont is essentially unaltered; higher values imply exposure to higher tem- peratures. Work by Bertram et al. ( 1992) and Cum- mins and Elderfield ( 1994) suggest that Sr isotopes are not altered in specimens with CAI-values of < 2.0. All samples examined in this study had CAI-values (deter- mined by B.R. Wardlaw) of 1, except for the Crittenden Springs, Nevada, specimens which had values of 2.

4. Methods

The conodonts were removed from the host rock by B.R. Wardlaw using standard procedures. The sur- rounding matrix was dissolved in 10% acetic acid, the

b. Detail of Permian/Triassic *‘S#‘%rSr data vs. age for samples from this study which define the rapidly decreasing and increasing segments of the curve. Symbols: U.S.A. -Texas, A = SR, V = ST, D = TX, U = BM; Wyoming, d = BC; Nevada, 0 = CS; Pakistan •- O=SA. ??=KL,O=CD, ??=KA. @=CW, M =NN.

c. The modified 87Sr/86Sr vs. age data set used to calculate the smoothed spline represented by the line. Data points enclosed in a dashed circle qpear to have been altered under continental condi- tions based on a comparison to our data as well as data from the literature, therefore they were excluded from the data set prior to curve fitting. Outlying data points, delineated with a cross ( X ), were eliminated using a systematic method described in the text. Symbols: O-U.S.A.; ??=Paktstan; @ =U.S. data eliminated for spline fit; ??= Pakistan data eliminated for spline fit.

E.E. Martin, J.D. Macdougall /Chemical Geology 125 (1995) 73-99 81

Table 1 Sr isotope and concentration data for Permian/Triassic conodonts

Sample” Weight (WI

Biozoneb Stage AgeC (Ma)

“SrlwSp 20 SP (ppm)

U.S.A.:

Neva Limestone, Kansas:

KN-1 37 A KN-2 38 B

Garden Vallqy Formation, Nevada:

NV-I 65 C NV-2 12 D NV-3 19 E

Skinner Ranch, Glass Mountains, Texas:

Assel _ 294 0.708111 o.OOoO22 1,563 Assel -2% 0.708132 o.oOOO22 1,614

Assel -289 0.708878 o.OOOo22 2,244 Sak _ 286 0.708841 0.000022 1.832 Art -215 0.707957 0.000022 799

SR-1 16 G SR-2 6 I SR-3 5 J SR-4 7 K SR-5 28 L SR-6 68 L SR-7 20 L SR-8 31 M SR-9 57 M SR-10 73 M SR-11 45 M SR-12 36 M

Split Tank, Glass Mountains, Texas:

Art 267.10 0.708011 0.000022 1,293 Art 266.90 0.708093 0.000033 1,517 Kung 266.70 0.708345 0.000030 1,276 Kung 266.50 0.707605 0.000026 1,879 Kung 266.30 0.707453 O.OCOO24 1,186 Kung 266.10 0.707629 0.000022 1,750 Kung 265.90 0.707643 o.OOoO22 1,829 Kung 265.70 0.707561 o.OOoO22 1,904 Kung 265.50 0.707506 o.OoOO22 1,375 Kung 265.30 0.707513 0.000022 1,810 Kung 265.10 0.707416 0.000084 1,482 Kung 264.90 0.707461 0.000022 1,666

ST-l 20 M ST-2 19 N ST-3 66 N ST-4 + 5 17 N ST-6 20 P

Del Narte Mount&s, Texas:

Kung 264.75 0.707336 0.000022 1,642 Kung 264.60 0.707341 0.000022 1,625 Kung 264.40 0.707333 0.000022 1,344 Kung/Uf 264.00 0.707388 0.000022 1,522 Uf 262.10 0.70733 1 o.OOoO22 1,598

TX-2 17 TX-3 33 ; TX-5 24 Q TX-6 24 R

Bird Mine, Del Norte Mount&s, Texas:

KaZ 261.75 0.707229 o.OoOO22 1,742 KaZ 261.00 0.706914 0.000022 1,533 CaP 258.25 0.707031 0.000022 1,415 CaP 257.80 0.706948 0.000032 1,543

BM-1 12 Q BM-2 28 R BM-3 36 R BM-4 30 S BM-5 41 R

Baldwin Creek, Wyoming:

Cap 258.25 0.707247 GP 257.80 0.707406 CaP 256.30 0.707059 GP 256.10 0.707027 Cap 256.50 0.706913

BC-1 55 N Kung 264.50 0.707340 BC-2 25 N Kung 264.30 0.707404 BC-3 61 N Kung 264.10 0.707373 BC-4+5 63 o/u Knz/Uf 262.25 0.707133 BC-6 19 U KaZ 260.75 0.707208 BC-7 31 U KaZ 260.00 0.707100 BC-8 41 V KaZ 259.25 0.70709 1 BC-9 11 V KaZ 258.60 0.707063

0.000022 0.000022 o.OOoO22 0.000022 0.000022

o.OOOo22 0.000022 0.000022 0.000022 0.000022 0.000022 o.OoOO22 o.OoOO22

1,485 1,790 2,109 1,672 1,443

1,424 1,469 1,441 1,226 1,730 1,339 1,306 1,148

82

Table 1 (continued)

E.E. Martin, J.D. Macdougall /Chemical Geology 125 (1995) 73-99

Sample” Weight (pgLg)

Biozoneb Stage Age” (Ma)

?w =sP 20 Sf (ppm)

U.S.A. (cont.):

Crittenden Springs, Nellada:

cs-1 24 cs-2 68 cs-3 29 CM 29

PAKISTAN:

Saidu Wali, Khisor Range:

SA-1 f 2 34 SA-3+4 16 SA-5 + 6 14

Kotia Lodlian, Khisor Range:

KL- 1 36 KL-2+3 16 KL-4+5 20 KL-6 25 KL-7 11 KL-8 20

KL-9 23 KL- LO 3

Chhidru Nab, Salt Range:

CD-l +2 4 CD-3 6 CD-4 7 CD-5 11 CD-6 10 CD-7 31

Kafhwai, Salt Range:

KA-1 21 KA-2 25 KA-3 6 KA-4 13 KA-5 17

Chatuwala Nala, Salt Range:

cw-1 7 cw-2 10 cw-3 31 cw-4 54

Nammal Nala,‘Salt Range:

NN-1 57 NN-2 18 NN-3 62 NN-5 38 NN-6 58 NN-7 25

FF Gries 250.00 0.707657 o.OOOQ22 1,395 II Smith 248.50 0.707885 0.000022 1,631 II Smith 248.40 0.707867 0.000022 1,960 JJ Smith 248.25 0.707940 o.OOoO22 1,848

T/U V V

V KaZ 260.75 0.707190 0.000022 1,400 V Cap 257.25 0.707168 0.000022 1,748 Y Wuch 255.70 0.70720 1 0.000022 1,538 Y Wuch 255.25 0.707100 0.000022 1,326 Y Wuch 254.80 0.707102 0.000023 962 Y Wuch 254.60 0.707117 o.OoOO22 1,436 Z Wuch 254.10 0.707186 0.000022 1,450 AA Wuch 253.75 0.707240 0.000054 1,417

V Kaz/Cup 258.60 0.707130 AA Wuch 254.30 0.707399 cc Chang 253.40 0.707690 cc Chang 252.70 0.707322 DD Chang 251.80 0.707427 HH Dien 249.00 0.707736

cc Chang 252.30 0.707499 cc Chang 252.20 0.707568 DD Chang 252.00 0.707458 DD Chang 251.80 0.707587 DD Chung 251.70 0.707469

cc Chang 251.80 0.707440 DD Chang 251.60 0.707503 EE Gries 250.75 0.707308 FF Gries 249.75 0.707626

FF Gries 249.75 0.707856 GG Dien 249.40 0.707895 GG Dien 249.25 0.708197 HH Dien 248.75 0.708178 HH Dien 248.60 0.708377 JJ Smith 248.35 0.708203

KaZ 260.75 0.707211 0.000022 1,396 KaZ 259.30 0.707388 0.000022 2,091 KaZ 258.75 0.707117 0.000022 1,871

0.000022 0.000034 o.OOoO22 o.OOOo22 0.000022 0.000022

0.000062 0.000022 0.000022 0.000049 0.000022

0.000024 0.000030 0.000022 0.000022

o.OOoO22 0.000022 0.000022 o.OOOo22 0.000022 o.OOw22

2,035 1,891 1,515 1,404 1,164 1,400

1,466 1,431 1.3% 1,490 1,460

1,556 1,572 1,511 1,425

1,241 n.a. 1,400 1,404 1,611 1,223

E.E. Martin, J.D. Maca’ougall /Chemical Geology 125 (1995) 73-99 83

Table 1 (continued)

n.a. = samples that were not analyzed for Sr concentration. “Sample numbers assigned for this study. For correlation with original sample designations, please contact first author. %onodont zonation from Wardlaw (1995). A = S. waubaunsensis4 “barskooi”; B = S. "longissimus ‘SW. expansus; C = S. fusus-M. longifoliosa; D = SW. inornatus; E = M. bisselli-Sw. primus; F = M. bissel&Sw. behnkeni; G = M. bisselli-Sw.whirei; H = N. clarki; I = N. pequopensis-M. gujioensis; I = N. exsculpfus-M. gujoensis; K = N. foliatus; L = N. foliates-M. idahoensis; M = N. prayi-M. i&hoer&; N = N. sukopilcatus-M. idahoensis; 0 = N. newelli-M. serrata; P = M. serrata; Q = M. asserrata; R = M. postserrata; S = M. nuchalina-M. n. sp.; T = M. phosphoriensis; U =: M. bitteri; V = Me. diaergens; W = M. “postbitteri”; X = M. liangshanensis; Y = M. leveni; Z = M. guanyuanensis; AA = M. orientalis; BB =M. “chanxingensis”; CC = M. subcarinara; DD = M. carinata; EE = H. parvus; FF = I. isarcica; GG = Ne. cristagalli; HH = Ne. pakistanensis; II = Ne. waageni; JJ = N. milleri. ‘Ages reported for combined samples represent a weighted mean of individual samples. d s’Sr/s%r values for standards equal 0.710260 for NBS 987, and 0.709175 for seawater samples from the North Atlantic and Central Pacific. All ratios are fractionation corrected to an %3/ssSr ratio of 0.1194. All samples have been assigned a minimum 20 uncertainty of f 22. 10m6 equivalent to the total range of repeat NBS 987 analyses. ‘Sr concentrations were determined by isotope dilution.

residue was wet-sieved and conodonts were concen- trated from the 20 to 180 mesh fraction using heavy liquids (sodium polytungstate) and hand-picked. At Scripps Institution of’ Oceanography, specimens from each sample were examined using scanning electron microscopy (SEM) tlo identify possible sources of con- tamination such as crystal overgrowths. In a few sam- ples small amounts of matrix material still adhered to the specimens.

Each sample was leached for - 20 min in a sonic bath with 1.5 M acetic acid. This process was repeated until each sample 1o;st - 15% by weight. This proce- dure was selected based on leaching experiments car- ried out on four large specimens. These specimens were dissolved in five sequential steps in which the amounts dissolved were - lO%, - 20%, - 20% and - 25%. 87Sr/86Sr in the first leachate was significantly more radiogenic, and the second leachate was slightly more radiogenic, than succeeding leachates, while values for the third, forth and final leachates generally agreed within uncertainly. All specimens were re-examined using the SEM after leaching. In most cases this tech- nique removed any :remaining matrix material as well as the surface layer of the conodont, thereby exposing the crystallites which comprise the lamellae (Pietzner et al., 1968; Barnes et al., 1973). Bertramet al. ( 1992) also found that minor surface leaching improved the interspecific correlations of 87Sr/86Sr in conodonts.

Diagenetic alteration of Sr isotopes is always a potential problem, particularly for Paleozoic speci- mens. One of the best tests against this possibility is a comparison of the i;sotopic ratios from widely distrib- uted sites. Because seawater Sr isotopes are a global

signal, samples of the same age exposed to a variety of environmental and diagenetic conditions should yield the same isotopic ratio if they have not been altered. With the exception of a few outliers, our data from a range of sites within the U.S.A. and within Pakistan agree well (Fig. 3b), but more significantly there is excellent agreement between U.S. and Pakistan sam- ples (Fig. 3c) over the time interval of overlapping sections. In addition, our data generally agree with pub- lished data from other locations and, most importantly, from other mineralogies, such as bulk carbonate and low-Mg calcite brachiopod shells. These mineral phases have different diagenetic susceptibilities than apatite.

Following surface leaching, conodonts were dis- solved in 1.8 iVHC1, spiked for Sr and Sm/Nd concen- tration measurements, and dried. A sheath of organic matter commonly remained. This was oxidized by add- ing 25 ~1 each of concentrated HN03 and 9 N HCl. At this point the sample was again dried, then processed through cation-exchange column chemistry using stan- dard techniques to separate Sr, Nd and Sm. Elemental concentrations, and 87Sr/86Sr and ‘43Nd/‘*Nd were analyzed by isotope dilution and thermal ionization mass spectrometry (TIMS) . Blanks for this procedure are 15 pg for Sr, 13 pg for Nd and < 1 pg for Sm. Initially, Rb concentrations were also measured using isotope dilution in order to correct *‘Sr/%r values for in situ decay of Rb. However, Rb concentrations were very low ( < 2 ppm) and the correction, which ranged from O.OOOOO2 to 0.000007, was insignificant com- pared to analytical uncertainties. Thus the uncorrected data may be as much as O.OOOOO7 low.

84

Table 2

E. E. Martin, J. D. Macdougall / Chemical Geology 125 (1995) 73-99

Nd isotope and concentration data for Permian/Triassic conodonts

Sample Weight Age” Sm Nd ‘47Sm/‘44Nd “=Nd/‘“Nd” (0) ‘“Nd?NdC (0 2C+ zNd(0 (CLg) (Ma) (ppm) (ppm)

U.S.A.:

Skinner Ranch, Glass Mountains, Texas:

SR-1 16 267.10 18.04 SR-6 68 266.10 14.00 SR-7+8 51 265.80 21.85 SR-9 57 265.50 16.94 SR-10 73 265.30 8.06 SRI]+12 81 265.00 3.04

Split Tank, Glass Mountains, Texas:

ST-2 19 264.60 8.07 ST 3-5 83 264.30 7.71 ST-6 20 262.10 47.67

Del Norte Mountains, Texas:

TX l-3 78 261.60 11.40 TX 4-6 70 258.40 19.50

Bird Mine, Del Norte Mountains, Texas:

74.01 0.147 0.5 12060 0.511819 27 -9.7 62.20 0.137 0.512136 0.511912 38 -1.9 94.99 0.139 0.512206 0.511979 14 -6.6 77.69 0.132 0.512149 0.511933 14 - 7.5 41.50 0.117 0.512190 0.511998 18 -6.2 14.56 0.126 0.512191 0.511985 22 -6.5

38.40 0.127 0.512109 0.511901 68 -8.1 34.56 0.135 0.512139 0.511918 18 -7.8

128.80 0.224 0.5 12262 0.511895 14 - 8.2

44.70 0.154 0.512211 0.511959 14 -7.0 86.30 0.137 0.512216 0.511992 14 -6.3

BM-1 12 258.25 BM-2 28 257.80 BM-3 36 256.30 BM-4 30 256.10

Baldwin Creek, Wyoming:

BC1+2 80 264.45 BC3 61 264.10 BC4+5 63 262.25 BC-6,7,9 67 259.75

Crittenden Springs, Neoada:

cs-1 24 250.00 cs-2 68 248.75 cs-4 29 248.20

PAKISTAN:

Saidu Wali, Khisor Range:

SA l-6 64 260.00

Kota Ladlian, Khisor Range:

KL-1 36 260.75 KL 2-5 36 256.50 KL-6 25 255.25 KL-7 11 254.80 KL-8 20 254.60 KL-9 23 254.10

18.90 95.74 0.119 0.512171 0.511976 21 -6.6 10.84 48.10 0.136 0.512148 0.511925 27 -7.6 26.99 109.00 0.150 0.512134 0.511889 18 -8.3 29.75 116.90 0.154 0.512116 0.511864 27 -8.8

6.27 39.60 0.096 0.512301 0.512145 14 -3.3 2.47 14.41 0.104 0.5 12343 0.512173 19 -2.8 2.23 12.83 0.105 0.5 12208 0.512036 24 -5.5 4.81 29.95 0.097 0.5 12255 0.512096 21 -4.3

1.26 7.60 0.103 0.512071 0.511902 62 -8.1 3.04 17.66 0.100 0.512104 0.511940 68 -1.3 4.86 27.90 0.105 0.5 12092 0.511920 59 -7.1

44.19 182.00 0.147 0.5 12023 0.511783 14

37.90 164.00 0.139 95.94 373.00 0.155 84.63 366.90 0.139 43.42 194.70 0.135

141.40 618.00 0.138 117.00 733.00 0.146

KL-10 3 253.75 82.80 372.10 0.134

- 10.4

0.5 12050 0.511819 17 -9.7 0.5 12277 0.5 12023 14 -5.7 0.512212 0.511984 14 -6.5 0.512280 0.512059 18 -5.0 0.5 12273 0.5 12047 15 -5.2 0.5 12242 0.512003 14 -6.1 0.512196 0.511976 84 -6.6

E.E. Martin, J.D. Macdougall / Chemical Geology 125 (1995) 73-99 85

Table 2 (continued)

Sample Weight Age” Sm Nd ‘47Sm/‘*ONd ‘.“Nd/‘“Ndb (0) ‘43Nd/‘UNdc (0 28 G.UCN (CLg) (Ma) (ppm) (ppm)

PAKISTAN (cont.):

Chiddru Nala, Salt Range:

CD-3 6 2154.30 CD-4 I 2!53.40 CD-5 11 :!52.70 CD-6 10 :!5 1.80 CD-7 31 :!49.00

Kathwai, Salt Range:

149.30 631.00 0.143 0.512174 0.511940 25 -1.3 94.01 385.70 0.147 0.512122 0.511881 30 - 8.5 44.82 153.00 0.177 0.5 12038 0.511748 67 -11.1 94.42 434.00 0.131 0.5 12080 0.511865 25 -8.8

6.99 30.56 0.145 0.512133 0.511910 34 -7.9

KAl+2 46 252.25 KA-3 6 252.00 KA-4 13 251.80 KA-5 17 125 1.70

Chatuwala Nala, Salt Range:

56.12 264.00 0.128 0.512134 0.511925 14 - 7.6 45.80 168.30 0.164 0.5 12085 0.511816 50 -9.8 40.73 139.30 0.177 0.512140 0.511851 14 -9.1 22.93 95.05 0.146 0.512095 0.511856 18 -9.0

cw- 1 7 cw-2 10 cw-3 31 cw-4 54

Nammal Nab

251.80 82.20 347.20 0.143 0.512133 0.511899 21 -8.1 251.60 110.00 405.00 0.164 0.512164 0.511895 22 -8.2 250.75 30.25 116.90 0.156 0.512010 0.511754 19 -11.0 240.75 10.32 43.60 0.143 0.512169 0.511935 26 - 7.4

NN-1 57 249.75 22.95 101.60 0.136 0.512174 0.511951 19 -7.1 NN-2 18 249.40 17.60 82.34 0.129 0.512107 0.5118% 32 -8.2 NN-3 58 249.25 6.49 25.42 0.154 0.5 12053 0.511800 23 - 10.1 NN-5 44 248.75 17.25 62.06 0.168 0.5 12036 0.511761 21 - 10.8 NN-6 58 248.55 22.71 83.80 0.164 0.5 12022 0.511754 14 -11.0 NN-7 47 248.35 4.98 24.29 0.124 0.512117 0.511915 25 -7.8

“Ages reported for combined samples represent a weighted mean of individual samples. “Nd isotopic value measured from conodonts. The ‘43Nd/‘“Nd value for the La Jolla standard is 0.511859. All ratios are fractionation corrected to ‘“NdO/ ‘“NdO = 0.242436. ‘Nd isotopic values from the Permian and Triassic corrected for radiogenic production of ‘“Nd. ‘43Nd/‘UNdc,, = ‘43Nd/‘UNdC,,, - 14’Sm/ ‘“Nd[exp(ht) - I]. “Uncertainty X 106. All samples have been assigned a minimum uncertainty of f 14. 10e6, equivalent to the total variability of repeat standard analyses. ebNd(r) = [‘43Nd/‘44Nd~,)I’143Nd/‘~NdlCHUR) - 1] X l@. ‘43Nd/‘UNdo,,,(,, ~0.512638. ‘43Nd/‘C(Nd(CHUR)(2MM.) =0.512316.

All samples were analyzed on a W-54@ single-col- lector TIMS at Scripps Institution of Oceanography. Final weights for samples composed of l-5 conodont elements, or pieces of elements, ranged from 4 to 70 pg. Sr concentrations were 1600 ppm *40% (Table 1). For isotope ratio measurements, sample amounts of - 25 ng Sr were loaded on tantalum oxide on a single tungsten filament and 300 ratios were col- lected at 1.5 V. The measured value and total variability for *‘Sr/*‘%r in 7 1 analyses of NBS 987 Sr loaded with this technique and analyzed over the interval of this

project was 0.7 10260 f 0.000022. This external preci- sion represents the minimum uncertainty assigned to any sample. For smaller samples we used a variation of this loading technique described by Birck ( 1986). For lo-ng standards the absolute value and precision using this small sample method is identical to that of the tantalum oxide technique; however, for very small samples the within-run uncertainty was often higher (Table 1).

Nd concentrations were highly variable, ranging from 8 to 600 ppm (Table 2). By analogy to fish teeth,

86 E.E. Martin, J.D. Macdougall /Chemical Geology I25 (1995) 73-99

rare-earth element (REE) uptake probably occurred after deposition of the conodont on the sea floor. Var- iations in concentration, therefore, may relate to expo- sure time and redox conditions in this environment. As might be expected, Sm/Nd ratios are far more consis- tent. At the beginning of the project the REE fraction of several samples were combined prior to the Sm-Nd separation to insure sufficient Nd was present for iso- tope ratio analysis. This practice was eventually dis- continued as knowledge of the expected concentration from each sample and the chemical yield improved. Nd samples were loaded on a single Re filament and ana- lyzed as NdO+ at 0.7 V for 300 ratios. Lower intensities were used and fewer ratios measured for smaller sam- ples. For 19 analyses of the La Jolla Nd standard ana- lyzed during this project the measured e,,-value was - 15.2 with a total range of k 0.27. Again, the within- run uncertainty for samples containing very small amounts of Nd (OS-2 ng) was generally higher (Table 2).

5. Discussion

5.1. Sr isotope results

The Pakistani and U.S.A. data outline a very rapid decrease followed by an even more rapid increase in seawater 87Sr/86Sr in the Late Permian (Fig. 3; Table 1) . Prior to the late Early Permian the distribu- tion of data points is too sparse to define precisely the shape of the isotope curve (Fig. 3a). A greater sample density from the Artinskian/Kungurian boundary to the Capitanian details s’Sr/%r values which decrease at an average rate of 0.000062 Ma- ’ to a minimum value of 0.70706 (Fig. 3b and c) . This rate of change exceeds the average rate of increase for the past 40 Ma (DePaolo and Ingram, 1985; Hess et al., 1986; Hodell et al., 1989, 1990) and the minimum represents one of the lowest points during the Phanerozoic with the pos- sible exception of the Jurassic (Burke et al., 1982). From the Capitanian through the P/Tr boundary sea- water ratios increase at the very rapid rate of O.oooo97 Ma-‘, approximately two and a half times the average rate of increase for the past 40 Ma. Based on data that continue beyond the earliest Triassic, this increase per- sists into the Anisian stage of the Triassic (Burke et al., 1982; Koepnick et al., 1990).

Veizer and Compston ( 1974)) Popp et al. ( 1986b), Koepnick et al. ( 1990), Kramm and Wedepohl ( 1991), Nishioka et al. ( 1991), and Denison et al. ( 1994) have all published 87Sr/86Sr data for this time interval. Because of differences in the time scales used for the individual data sets, interlaboratory correlations are exceedingly difficult. In addition, the method of stratigraphic correlation is not always defined in other studies, and it is likely it differs from the conodont biostratigraphy applied to our data. Veizer and Comp- ston ( 1974) and Popp et al. ( 1986b) assigned sample ages based on sample location in the upper or lower portions of a stage. In these cases, we assumed that their stages are consistent with those utilized in this study. Nishioka et al. ( 1991) used a very similar time scale, thus we have used their age estimates shifted by 6 Ma to account for the revised P/Tr boundary age, Denison et al. ( 1994) used the time scale of Harland et al. ( 1990)) therefore their data have also been shifted by 6 Ma and those from the Ufimian through Artinskian adjusted slightly ( < 1 Ma). Kramm and Wedepohl ( 1991) only place their samples within the Zechstein stratigraphy, therefore we were unable to correlate their data exactly. By comparison to our 87Sr/86Sr values their data appear to be from the Late Kazanian to Changsingian, which is roughly equivalent to most esti- mates for Zechstein ages.

As Fig. 3a illustrates, some of the data points from Veizer and Compston ( 1974) and Poppet al. ( 1986b) and Nishioka et al. ( 1991) agree well with our values, while others are much lower. The data recently pub- lished by Denison et al. (1994) are the only other set with sufficient coverage to delineate the curve. The correlation between their data and ours is impressive, although their values tend to be lower in the latest Permian, possibly due to differences in correlation techniques or the time scales applied. As a conse- quence, the 87Sr/86Sr minimum appears to occur at a slightly younger age in their data set than in ours.

As mentioned earlier, the calculated rate of change of seawater 87Sr/86Sr is highly dependent on the time scale employed. The numeric ages from Odin (1982) or Forster and Wanington (1985) would produce a more gradual change for the decreasing portion of the curve than that in Fig. 3; however, the rate of increase in the latest Permian would be _ 1.5 times more rapid. Little change would result if the time scale of Palmer ( 1983) were used. Middle and Late Permian stages are

E. E. Martin, J.D. Macdougall /Chemical Geology 125 (I 995) 73-99 87

much longer in the time scale proposed by Menning ( 1989) than in Harland et al. ( 1990), thus both the decrease and increase :in isotopic ratios would be more gradual.

5.2. Sr geochemical model

Using equations that describe the Sr geochemical cycle it is possible to estimate the magnitude and pos- sible causes of change.s in the Sr budget that would be required to produce the observed seawater 87Sr/86Sr fluctuations during the Middle Permian to Early Tri- assic. There are three major sources of Sr to the oceans: river water, which transports weathered continental material to the oceans; hydrothermal fluids, which have interacted with mid-ocean ridge basalts at high tem- peratures; and sediment pore fluids, composed of sea- water that has been altered by carbonate dissolution and recrystallization at low temperatures. According to Palmer and Edmond ( 1989), the modem 87Sr/86Sr ratios for these sources are estimated to be 0.7119, 0.7035 and 0.7084, respectively (the pore fluid value was taken from Elderfield and Gieskes ( 1982) ) , while the relative proportions of each are N 65%, * 30% and * 5%.

The rate of change of seawater 87Sr/86Sr is described by the following equation:

&w ~=fwW,(R”-~sw)l (1)

where N is the number of moles of Sr in the oceans; J, is the flux of Sr (in mol a- ’ ) into the ocean from source n; R, is the “Sr18?Sr ratio of source n; and Rsw is the seawater 87Sr/86Sr ratio (Hodell et al., 1989; Kump, 1989; Capo and DeP,aolo, 1990; Richter et al., 1992).

As Kump ( 1989) pointed out, the pore-water flux is small and its isotop:ic ratio approaches the seawater value; thus, this source can be neglected without sig- nificantly affecting the model. Richter et al. ( 1992) estimated that ignoring this flux introduced an uncer- tainty of N f 5% in their calculations of other fluxes. We also ignore this term in the calculations. Thus Eq. 1 becomes:

d&w N -==J,(R,-Rsw) +Jt,(&-Rsw) dt (2)

where subscripts r and h symbolize the riverine and hydrothermal inputs, respectively.

We test the sensitivity of the model to changes in the various unknowns in order to determine which values are most appropriate for the Permian, and to evaluate the potential of each parameter to control seawater 87Sr/86Sr variations. Of the seven variables incorpo- rated in this equation, the seawater 87Sr/S6Sr (R,,, and the rate of isotopic change (d&,ldt) are based on analytical measurements, leaving five unknowns. As in previous models (Kump, 1989; Capo and DePaolo, 1990; Richter et al., 1992) the 87Sr/86Sr ratio of the hydrothermal component (R,,) is considered to be sim- ilar to the mid-ocean ridge basalt value and constant through time. We have chosen a value of 0.703, which is slightly less radiogenic than the ratio given by Palmer and Edmond ( 1989), but it is probably a more accurate reflection of the end-member hydrothermal contribu- tion (Richter et al., 1992). This assumption further reduces the number of unknowns to four. For the Ceno- zoic portion of the curve Richter et al. ( 1992) elimi- nated an additional variable by assuming that the hydrothermal flux was proportional to sea-floor spread- ing rates; however, these are unconstrained for the Per- mian. For relatively recent times it may be safe to assume that the modem estimates of R, and J,, are rea- sonable approximations of the true value, but this is much less certain for the Permian. As already dis- cussed, the combination of continental configuration, sea level, and resulting climate conditions were prob- ably quite unique for this period. These factors would likely influence weathering and drainage patterns as well as hydrothermal circulation, which in turn could generate large changes in the geochemical cycle of Sr. However, as discussed below, it may be possible to constrain changes in the Sr isotopic ratio of the riverine flux (R,) by examination of seawater Nd isotopes.

Input values for Rsw and dR,,ldt were calculated from a smoothed spline fit to a modified data set (Fig. 3c), with outliers eliminated from the data prior to curve fitting. First the three oldest samples from Texas were removed because of their anomalously radiogenic ratios compared to all other data (Fig. 3a). The increasing and decreasing segments of the seawater curve were each then fit with a second-order polyno- mial. Any isotopic ratio which deviated more than 0.00025 ( -0.035%, or ten times analytical uncer- tainty) from these curves was removed, and new curves were fit to the remaining data. This procedure was repeated a second time, eliminating values which devi-

88 E.E. Martin, J.D. Macdougall /Chemical Geology 125 (1995) 73-99

ated by more than 0.00015 (0.02%). This iterative smoothing process eliminated a total of four out of thirty-four data points ( 12%) from the decreasing seg- ment, and seven out of twenty-eight (22%) from the increasing segment; however, four of those were from a single Triassic locality (Nammal Nala) .

Earlier works (e.g., Burke et al., 1982) fit curves to the minimum value measured for any given time inter- val, based on the assumption that diagenetic alteration in a continental setting would tend to increase the iso- topic ratio. However, in Fig. 3a it is clear that some of the ratios reported in this study, as well as some from earlier studies, are anomalously low relative to other samples of the same age that exhibit a variety of min- eralogies and come from widely distributed locations. These anomalous samples probably do not represent the best-estimate of seawater in the past. However, our spline fitting technique did eliminate twelve points from above the curve but only two from below (Fig. 3c), suggesting that most anomalous samples were indeed ones for which 87Sr/86Sr has increased. We note also that a curve fit through the lowest values would have slightly different absolute values, but essentially the same shape as our spline-fit curve, and therefore would not significantly alter our model cal- culations.

5.3. Nd isotope results

Like Sr, the Nd isotopic composition of seawater potentially records changes in continental weathering; however, there are two important differences. First, there is no balancing input from sea-floor hydrothermal systems (Michard et al., 1983; Piepgras and Wasser- burg, 1985; Bertram and Elderfield, 1993). Instead, input to the ocean comes entirely from Nd dissolved in river water and adsorbed on river-borne and eolian particulates (Goldstein and Jacobsen, 1987). Second, the residence time for Nd in seawater is less than the mixing time of the oceans (Elderlield and Greaves, 1982; Piepgras and Wasserburg, 1987; Jeandel and Peng, 1989). As a result the ‘“Nd/‘#Nd ratios of individual water masses are distinct, reflecting the lithology and age of material weathered from the sur- rounding continents.

Nd isotopic compositions were determined for a selected group of the samples analyzed for Sr. These results are presented in Table 2 and Fig. 4, and con-

-12 ““““““““1 .‘I”” Triassic 1 LatePenian Middle Permian 1 Early Ch 1 WU ) Ca ) Ka 1 Uf 1 KU 1 A,

245 250 255 260 265 270

Age (Ma)

(b) -2~ .I ‘1 ‘. 1’. ‘. 1 “, 1, ” I

-4 -

s -6 -

r” w

-0 - u

0

-10 -

-12 ’ . ’ ’ ’ Triassic

Late Permian 1 Middle Permian 1 Early Ch 1 WU 1 Ca 1 Ka 1 Uf 1 KU 1 Ar

245 250 255 260

Age (Ma)

265 270

Fig. 4. Permian/Triassic eNd data vs. age for samples from: (a) the US; and (b) Pakistan. The arrows highlight decreasing Nd isotope trends indicative of increased continental input. For data points with- out error bars the error is less than the size of the symbol. Symbols as in Fig. 3.

stitute one of the first detailed records of short-term ‘43Nd/‘44Nd variations in seawater. Changes in oceanic Nd isotopic composition are considerably less system- atic than those of Sr. This probably reflects the short residence time, and therefore more rapid response, of seawater Nd to changes in the inputs, as well as varia- bility due to local rather than global phenomena. This point is illustrated by the offset between Wyoming and Texas samples (Fig. 4a). Both of these locations were situated along the PaleoPacific margin during the Per- mian, but their distinct Nd isotopic ratios imply that mixing between these sites was incomplete. Another possible explanation for the scatter could be diagenetic alteration, although, as we have argued earlier, the low CAI for these samples suggests that this problem should be minimal

E. E. Martin, J.D. Macdougall / Chemical Geology 125 (1995) 73-99 89

As is the case for modern phosphatic fish teeth, con- odonts probably contained only ppb levels of Nd when the animal was alive, and the current high levels of Nd were incorporated following deposition on the sea floor (Wright et al., 1984; Shaw and Wasserburg, 1985). To dispute the claim that this “excess’ ’ Nd representspore fluids altered by surmunding sediment Staudigel et al. ( 1985) pointed out that fish teeth yield the same Nd isotopic values as Mn nodules which are known to form at the sediment-water interface in direct contact with seawater. Several authors have demonstrated that ratios for Mn nodules closely track seawater values and are distinct from surrounding pelagic clays (Piepgras et al., 1979; Elderfield et al., 1981; Goldstein and O’Nions, 1981). In fact, Piepgras et al. ( 1979) dem- onstrated that the Nd isotopic ratios from the top of a Mn nodule, which was in contact with seawater, and the bottom of the nodule, which was in contact with sediments, were identical, although the concentration of Nd on the top was twice that of the bottom. This suggests that seawater is the ultimate source of Nd to the fish teeth, and paesumably therefore to conodonts as well. There is also concern that phosphates continue to take up and exchange Nd during burial. To address this problem Bernat ( 1974), Staudigel et al. ( 1985), Elderfield and Pagett ( 1986) and Wright et al. ( 1987) all showed that recent fish teeth acquire very high con- centrations of Nd wil;hin the top few mm’s of the sed- iment-water interface, and that there is no systematic increase in this concentration with burial depth. It is possible that much of this early uptake coincides with alteration of the hydroxyapatite from the living fish to the carbonate apatite of the fossil specimen; although this association has yet to be studied in detail.

For our samples the most intriguing feature in the Nd data is the dramatic decrease to more continental values that occurred in the Late Permian in both the western U.S.A. (Paleo-Pacific) and Pakistan (Teth- yan) samples. This gross trend is apparent at every site evaluated from 260 to 250 Ma regardless of the asso- ciated Nd concentration. Concentrations in the U.S.A. samples are generally < 100 ppm Nd, while Pakistan samples range from 100 to > 700 ppm Nd. It has been suggested the Nd content of phosphates may be related to their alteration history. In a study of REE patterns preserved in fish teeth, Elderfield and Pagett ( 1986) concluded that specimens from oxic, slow sedimenta- tion rate, deep-sea sediments that had high Nd concen-

trations were the most likely to represent seawater chemistry accurately. In contrast, Bertramet al. ( 1992) found a correlation between Nd concentration and more continental Sr and Nd isotopic signatures in conodonts, such that specimens with higher concentrations were more likely to be altered. We found no correlation between Nd concentration and the deviation in 87Sr/ ‘?jr from the spline fit for our seawater curve, and conclude that for those samples there is no clear alter- ation signal for Nd that is related to concentration.

Assuming that the decreasing Nd values do reflect changes in seawater chemistry, these data help con- strain the possible causes of change in the Sr geochem- ical cycle. The Nd data indicate that the types of rocks weathered from the continent were increasingly “con- tinental”, which would also affect the 87Sr/86Sr values. In both locations the decrease in ‘43Nd/‘44Nd appears to begin close to the minimum in the Sr curve, although there may be small differences in timing between the two groups of samples.

5.4. Sr model sensitivity and results

5.4.1. Sr content of seawater We first examined the influence of oceanic Sr content

on other model parameters. Currently the number of moles of Sr in the ocean (N) is 1.19~10”. As men- tioned earlier, Holser and Magaritz (1987) have sug- gested that this value may have been lower in the Permian. Our calculations indicate that evaporite for- mation would have decreased seawater Sr concentra- tions less than w 6%.

In Fig. 5a we show the influence of N on the riverine flux required to produce the observed *‘Sr/?$r curve. Calculations are shown for three values of N, ranging from 30% below to 15% above the modem Sr content of the ocean. Although somewhat arbitrary, these val- ues encompass a total variation of almost 45% in N, and are based on the sum of the inputs required to satisfy the observed Permian seawater 87Sr/86Sr at the begin- ning of the model (265 Ma), divided by two estimates of the residence time for Sr in the oceans: 2.5 and 4.0 Ma. Although a modem residence time of 4 Ma is frequently cited (summarized in Elderfield, 1986), improved data on the total Sr fluxes to the oceans (Pal- mer and Edmond, 1989) yield a value closer to 2.5 Ma (Hodell et al., 1990). The model results were calcu- lated for each value of N using: (1) a modem hydro-

90 E. E. Martin, J.D. MacdougaN/ Chemical Geology 12.5 (199s) 73-99

2.50E+iO

4 E 3j 2.OOE+lO

e ‘C u

g 1.50E+iO

Fig. 5a

l.OOE+lO . . ’ . . ’ ’ ’ . ’ . . ’ . Triassic L&Permian I MiddlePermian I Early

Ch 1 Wu 1 Ca 1 Ka ] Uf ] Ku 1 A, 245 250 255 260

Age (Ma)

265 270 245 250

O.OOE+OO ’ ’ ‘,’ ’ ’ ‘,’ & Triassic

Late Permian Middle Permian EN Ch [ Wu 1 Ca 1 Ka 1 LH 1 Ku 1 Ar

255 260

Age (Ma)

266 270

(d) 4.50E+lO , . . , . . , . . , . .

3.50E+lO -

Modern Value

2.50E+lO -

1.50E+lO

t Fig. 5d

0.706 ’ ‘~t~p:rr;lian ’ Lddle ~,,i’,’ t’eri 050E+lO ....‘..,,‘....‘..,.‘,.,, Triassic Ch ( Wu ) Cs 1 Ka 1 Uf 1 Ku 1 Ar TriZSSiC

LatePermian 1 Middle Permian 1 Earfy Ch 1 WU 1 Ca 1 Ka 1 Uf I Ku ( Ar

245 250 255 260

Age (MaI

265 270 245 250 255 260

Age (Ma)

265 270

Fig. 5. Permian/Triassic “‘!W8%r model results for: (a) the calculated riverine flux (Jr) given a range of values for the number of moles of Sr in the ocean (N), and assuming J,, = 1.47-10” mol a-’ Sr, R, =0.703 and R, =0.7110; (b) the calculated hydrothermal flux (J,,) given a range of values for the riverine flux (Jr). and assuming N= 1.19-10” mol Sr, R,, =0.703 and R, =0.7110; (c) the calculated riverine isotopic ratio (R,) given a range of values for the riverine flux (J,). and assuming N= 1.19-10” mol Sr, J,, = 1.47-10” and R,,=O.703; and (d) the calculated river-me flux (Jr) given a range of values for the river-me isotopic ratio (R,) , and assuming N= 1.19 * 10” mol Sr, J,, = 1.47 - 10” and R,, =0.703.

thermal flux of 1.47 ??10” mol a- ’ Sr, which is slightly less than the value given by Palmer and Edmond ( 1989) to compensate for the fact that we ignore the pore-water flux; (2) an 87Sr/86Sr of 0.703 for the hydrothermal flux; (3) a riverine 87Sr/86Sr value of 0.7 110 which approximates the modern ratio (0.7 119) minus the highly radiogenic values for rivers draining the Himalaya; and (4) an assumption of steady-state conditions.

The results indicate that the model is not particularly sensitive to changes in N. A 45% change in N would be compensated by a maximum change of only 7% in the riverine flux (Fig. 5a) ; thus a 6% decrease in sea- water Sr concentration due to evaporite formation

would have a negligible effect on other model results. Therefore, in further calculations, we assume the mod- em value of 1.19*10” mol Sr for N.

5.4.2. Hydrothermalflux In Fig. 5b we plot the hydrothermal flux that would

be required to match the observed seawater curve if riverine flux values were held constant. Three different rivet-me fluxes are used: the modem flux of 3.330 10” mol a- ’ (Palmer and Edmond, 1989), as well as two lower values ranging down to approximately half the modern flux. The lower two values were calculated to yield the observed Rsw at the start of the model (265 Ma) and at the 87Sr/86Sr minimum (257.25 Ma)

E.E. Martin, J.D. Macdougall /Chemical Geology 125 (1995) 73-99 91

assuming the modern value for J,,, 0.703 for R,,, and 0.7 110 for R,.

As mentioned previously, the hydrothermal flux is probably directly related to mid-ocean ridge volume. Complete assembly of Pangea at the end of the Early Permian (Scotese and McKerrow, 1990) suggests ridge volume was probably at a minimum at this time. Based on the extremely low sea level this volume prob- ably remained low into the Late Permian. Thus the hydrothermal flux was likely to be less in the Permian than today, requiring a smaller balancing Permian riv- erine flux than the modern value (Fig. Sb) .

It can also be inferred from Fig. 5b that changes in Jh alone are not likely to have generated the observed seawater 87Sr/86Sr variations. Even using the lower riverine flux values, the hydrothermal flux would have to have increased by N 15% during the Middle Per- mian, then decreased by N 50% from the Late Permian into the Triassic. However, based on the distribution of ridges, in particular the initial rifting of South China in the Late Permian, thle hydrothermal flux should have increased throughout the time interval investigated. Even if the sense of change were correct, the Late Permian-Triassic change is unrealistically large, requiring a 50% decrease in ridge volume over only 10 Ma if the hydrothermal flux is directly proportional to ridge volume. This compares to a predicted maximum change of 10% per 10 Ma calculated for a single ridge system in the Cenozoic (Kominz, 1984).

X4.3. Riverine flux The riverine Sr flux required to explain the observed

seawater isotopic ratios is strongly anticorrelated with its *‘Sr/*?Sr ratio. Fig. 5c shows the riverine isotopic ratios required for thle three riverine flux values used in Fig. 5b, again assuming modern values for N, J,, and Rh. The range of isotopic compositions fall easily within the range reported for individual modern river systems (Goldstein and Jacobsen, 1987; Palmer and Edmond, 1989). However, the model results are for global average inpms. Note that for the two lower J,- values ( -45% and - 55% of the modem flux) high input *‘Sr/s”Sr values are required for much of the Late Permian and Early Triassic, We have argued above that J, in the Permian was probably lower than today; thus possible reasons for high riverine 87Sr/86Sr need to be assessed.

As mentioned earlier, decreasing seawater Nd iso- topes suggest the type of rocks weathered from the continents were becoming more “continental” during the Late Permian. Using our Nd data and the relation- ship between Sr and Nd isotopic ratios of dissolved river runoff published by Goldstein and Jacobsen ( 1987) yields a shift in 87Sr/86Sr values of 0.7091 to 0.7104 for 259 to 255 Ma from the U.S. data, and 0.7082 to 0.7104 for 255 to 250 Ma from the Pakistan data. Comparison of the U.S.A. results and Fig. 5c illustrates that given the modem riverine flux, the cal- culated change in the riverine isotopic ratio could account for the entire increase in seawater *‘Sr/*?Sr. However, the relationship proposed by Goldstein and Jacobsen ( 1987) is based on data from rivers, not sea- water. Applied to the present-day ocean, it would pre- dict *‘Sr/*‘?Sr values that are low compared to those observed for seawater or global runoff. This may in part reflect the different behavior of these two elements in seawater, especially the difference in residence time, and also the fact that the particulate Nd flux is not accounted for. The geochemical cycle for Nd is still poorly understood; for example, estuarine processes apparently strongly impact the relationship between riverine and seawater Nd isotopic compositions, but the effect has not been quantified. For our Permian samples, the Nd data provide information on only two regions, but seawater *‘Sr/*?Sr depends on knowledge of global inputs. Therefore, although the Nd data strongly suggest that riverine *‘Sr/?Sr increased in the Late Permian, it provides little information on the mag- nitude of that change. For the decreasing portion of the seawater *‘Sr/*?Sr curve, in the Middle Permian, Nd data are too few and too scattered to help constrain the Sr input.

The two most likely reasons for rapidly increasing riverine *‘Sr/@Sr are continental glaciations and Hi- malayan-style collisional events. Glacial activity can remove surface sediment and produce rock flour from old shield material, thereby greatly increasing the sur- face area of material with high 87Sr/86Sr exposed to weathering (Armstrong, 197 1; Palmer and Elderiield, 1985; Miller et al., 1991; Zachos, 1993). However, the only known glacial episode during the interval we investigated is in the Early Permian and coincides with decreasing Sr isotope ratios. There is no evidence of glacial activity during the Late Permian increase in 87Sr/86Sr.

92 E.E. Martin, J.D. Macdougall/Chemical Geology 125 (1995) 73-99

Edmond ( 1992) argued that it is difficult to alter the global riverine isotopic value substantially because of the inverse relationship between Sr concentration and 87Sr/86Sr in most drainage basins (Palmer and Edmond, 1989; Palmer and Edmond, 1992). However, modem rivers draining the Himalaya pose an exception to this relationship (Palmer and Edmond, 1989; Krish- naswami et al., 1992), which Edmond ( 1992) attrib- utes to the formation of high-grade metamorphic rocks during the continental collision, and the associated redistribution of radiogenic Sr into phases more sus- ceptible to chemical weathering. But available evi- dence concerning the sequence of tectonic events in the Paleozoic leading to the formation of Pangea indicates that the major continental collisions, and the erosion from these events, preceded the observed seawater 87Sr/86Sr increase by tens of millions of years. The youngest of these collisions resulted in the formation of the Urals; however, this mountain range was not located in the tropics where it would be most suscep- tible to weathering. Small Asian microcontinents and arcs may have collided in the Late Permian (Nie et al., 1990), but these would have lacked the radiogenic material necessary to alter the global isotopic ratio of river runoff. Thus, the source for increasing riverine “Sr/‘“Sr suggested by decreasing eNd- values is diffi- cult to identify. For this reason we believe it is unlikely that the increase was dramatic enough to account for the entire increase in the seawater isotopic ratio.

In the section above, we have shown that if other parameters are held constant, it is unlikely that changes in R,, N or J,, by themselves can account for the observed large variations in seawater 87Sr/86Sr during the time period between N 257 to u 250 Ma, although an increase in R, was probably a contributing factor. This leads us to the conclusion that there must have been significant changes in the riverine Sr flux. Fig. 5d illustrates the required variations in this parameter for three different values for RP Again R,,, N and J,, were assigned their modem values. The R, estimates are based on present-day values: Goldstein and Jacobsen (1987) report 0.7101 as the weighted average of 13 large rivers today, and this value is similar to the “Sri ‘?jr values calculated from the Nd isotope data; Palmer and Edmond ( 1989) estimated that they surveyed 47% of the total river runoff to obtain a global average of 0.7 119; and Capo and DePaolo ( 1990) and Richter et al. (1992) chose the intermediate value of 0.7110,

which also approximates the isotopic ratio of modem runoff excluding rivers draining the Himalaya. It is obvious from Fig. 5d that the riverine Sr flux required to match the seawater Sr isotope data is quite sensitive to even relatively small changes in the “Sr/@?Sr of the river input. The total range in the riverine flux calcu- lated using the three isotopic estimates is l.l-lO1o- 4.1 ??10” mol a- ‘, similar to model results for the last 100 Ma (Richter et al., 1992). The highest riverine 87Sr/86Sr 0.7119, requires the least variation in the riverine Aux; however, this value is derived from the present, a time when seawater 87Sr/86Sr is at its highest value since the beginning of the Phanerozoic. As dis- cussed earlier, there is no apparent source for such radiogenic runoff during the Late Permian. The lowest value investigated, 0.7101, requires the greatest change in the riverine flux, reaching values that exceed the modem input. As we have argued earlier, based on the probable low hydrothermal flux in the Permian, riverine fluxes were likely to have been lower than the present day. Thus, a value for R, near0.7 110 seems reasonable. Assuming that the isotopic ratio of the flux also increased over this interval, the magnitude of increase in the riverine flux would be less than that illustrated in Fig. 5d.

To summarize, we have investigated the sensitivity of the various parameters affecting seawater *‘Sr/*‘?jr to change in each of the other parameters. We conclude that the major cause of decreasing seawater *‘Srjg6Sr in the Middle Permian was a decrease in the riverine flux, while the major factors leading to the large changes in the seawater ratio in the Late Permian-Early Triassic were probably an increase in the riverine iso- topic ratio and the riverine Sr flux. The intermediate curve in Fig. 5d illustrates a doubling of the riverine flux over 12 Ma from 1.45*10” mol a-’ Sr in the Middle Permian to 2.90 10” mol a-’ Sr in the Early Triassic. This represents the maximum potential change required in the riverine flux. The magnitude would be significantly reduced by an increase in R,, the probably smaller hydrothermal flux, and a lower oce- anic Sr content due to evaporite deposition.

5.5. Implications of the model results

The model results (Fig. 5d) thus suggest that the riverine Sr flux decreased by _ 10-152 during the Middle Permian. This does not define the full extent of

E. E. Martin, J.D. Macdougall / Chemical Geology 125 (I 995) 73-99 93

the decrease because the seawater “Srls’?jr ratio actu- ally began decreasing earlier in the Early Permian (Fig. 3a). Geologic evidence suggests that the Her- cynian and Uralian orogens were still topographic highs subject to extensive erosion in the Early to Middle Permian. Thus the decrease in seawater 87Sr/86Sr implies that either the rocks exposed to weathering had low 87Sr/86Sr ratios, or chemical weathering products were not transported to the ocean.

The decreasing seawater isotopic ratios could be attributed to extensive continentality and cool condi- tions during this interval. The vast continental interiors would have been very dry (Kutzbach and Gallimore, 1989; Crowley and N’orth, 1991) and the proportion of internal runoff compared to that reaching the oceans may have been very high. In addition, evidence of gla- ciations at high southlern latitudes at this time suggests that the global climate was probably cool and dry, resulting in less chemical erosion than at warmer times (Brady and Carroll, 1994), Another factor contributing to cooler conditions and less intense chemical weath- ering may have been a drawdown in atmospheric CO* caused by the initial fstages of intense weathering asso- ciated with the orogenies. As Caldeira (1992) sug- gested, it is difficult to sustain high rates of chemical weathering unless there are renewed sources of COZ in the atmosphere, such as volcanism or metamorphism.

Evidence of arid conditions includes the deposition of massive evaporites which characterize much of the period, but are particularly prominent in the Kungurian and Kazanian (Fig. 1). Additional evidence for decreased continental runoff comes from a study of nonluminescent Carboniferous to Kungurian age brachiopods. Delaney et al. (1989) report that Li/Ca ratios in these fossils are only -50% of the value observed for modlern and Devonian specimens (Fig. 1) , and speculate that this may be the result of lower Li concentrations in seawater due to a lower flux from the continents.

The transition from decreasing to increasing sea- water 87Sr/8”Sr (and thus, in our model, to an increas- ing riverine isotopic ratio and flux) occurred in the earliest part of the Late Permian. This is puzzling because there are several factors which might be expected to force a continued decrease in seawater 87Sr/86Sr at this time. These include: reduced rates of erosion of Pangean oaogenies, which by this time would have subdued topographic expressions; a lack of evi-

dence for glaciation; and climate models which suggest persistent arid conditions into the Late Permian due to the continental configuration.

As discussed, it is difficult to understand the source of increasing riverine 87Sr/86Sr* however Erwin’s (1993) analysis of conditions in the Late Permian poses a possible solution to the increase in the riverine flux. He suggested that atmospheric CO* levels may have been high at this time, resulting in global warming, which could lead to enhanced chemical weathering. More intense weathering might also result in weather- ing of deeper erogenic roots, particularly in the Her- cynian mountains located in the tropics, and therefore might also explain the increase in riverine 87Sr/86Sr. Although there is no direct record of paleo-CO2 con- centrations, Erwin ( 1993) has compiled a convincing list of possible Late Permian CO:! sources, including release of methane from gas hydrates, oxidation of organic material, and emissions from volcanism. As sea level fell, large areas that previously lay within the hydrate window (300-500 m below sea floor) along the continental shelves and within intracratonic basins would have been exposed, creating unstable conditions for any hydrates. The methane released through this process is an effective greenhouse gas, and it converts to CO, over short time scales in the atmosphere.

Carbon isotope data provide evidence of another possible source of atmospheric CO*. As mentioned ear- lier, Si3C-values were quite high ( N 3-4%0) during the Early and Middle Permian, then decreased to values less than zero during the Late Permian (Fig. 1) (Popp et al., 1986a; Holser and Magaritz, 1987; Magaritz et al., 1988; Baud et al., 1989; Gruszczynski et al., 1989; Holser et al., 1989). This decrease may in part reflect input of isotopically light carbon from release of gas hydrates. Another viable source of light carbon would be oxidation of organic carbon associated with falling sea level and exposure of shallow shelves (Holser et al., 1989; Erwin, 1993; Grossman, 1994). The net result of these processes would be a decrease in the organic carbon reservoir.

Large amounts of CO2 would also have been released during the eruption of the Siberian Traps, which have an estimated volume of > 1.5 - lo6 km3 of basalt. The amount of CO1 released during this eruption could be as much as lOI mol, or lo’* mol a-’ if spread evenly over 1 Ma (Erwin, 1993 -extrapolated from McLean, 1985). For comparison, based on Bemer’s ( 1990,

94 E.E. Martin, J.D. Macdougall /Chemical Geology I25 (1995) 73-99

199 1) model, the Permian atmospheric CO* reservoir was two to four times greater than the modern reservoir, on the order of 10L6-10” mol. Although the Siberian Traps appear to have erupted after the minimum in s7Sr/‘%r and the total output of CO* was relatively small, they may have provided a source of CO* to sustain enhanced weathering rates as CO* liberated from other sources was consumed.

Berner ( 1990, , 1991) has modeled atmospheric CO* levels in the Permian based on a set of criteria distinct from those considered by Erwin ( 1993). Ber- ner’s model is based on a combination of estimates of volcanic and metamorphic degassing derived from sea- floor spreading rates (ultimately from sea level), sed- iment burial rates estimated from isotopic data, weathering rates calculated from the distribution, area and elevation of the continents, land plant evolution, and estimates of river runoff. He found low levels of CO* in the Early Permian, similar to modern values, but increasing concentrations into the Triassic, which supports Erwin’s ( 1993) proposal.

A warmer, more humid climate in the Late Permian, initiated by increasing atmospheric CO*, is also con- sistent with mid-latitude lithologic, fauna1 and floral distributions. Yemane (1993) studied a region in Africa located between 45’ and 60”s during the Late Permian, for which climate models predict extreme seasonal temperatures and arid conditions. Yet the geo- logic record is one of extensive lacustrine deposits, therapsid fossils and Glossopteris vegetation, suggest- ing humid, temperate conditions. Taylor et al. ( 1992) also describe a Late Permian fossil forest from a paleo- latitude of W-85%. The fossils represent a deciduous forest in which widely spaced growth rings and the absence of frost rings indicate warm conditions at high latitudes.

Finally, decreasing oxygen isotopes in the latest Per- mian of the Carnic Alps, Austria (Holser et al., 1989, 1991) support the idea of greenhouse warming. Although this decrease has not been identified globally and it could possibly be attributed to diagenesis or salinity changes, Holser et al. (1991) argue that it rep- resents a - 6°C rise in the temperature of seawater.

Currently most of the chemical weathering of the continents occurs in a band surrounding the equator. If this situation were also true for the Permian, orographic effects of the low-latitude mountains generated by the Hercynian megasuture may have had a significant

impact on chemical weathering at that time. During the Late Carboniferous/Early Permian these mountains may have cast a rain shadow over Pangea (Scotese and McKerrow, 1990). As they eroded, the rain-laden equatorial easterlies would have penetrated farther onto the continent. Parrish (1993) has also suggested that the intensity of monsoonal circulation (Robinson, 1973; Kutzbach and Gallimore, 1989), and therefore the hydrologic cycle, increased throughout the Permian to a maximum in the Triassic. Thus erosion of core regions of equatorial mountains coupled with the devel- opment of monsoonal circulation may have accentu- ated the increase in chemical weathering, and thus the riverine Sr isotopic ratio and flux, associated with increased levels of atmospheric COZ.

Deep-sea sediment accumulation rates and the depth of the carbonate compensation depth (CCD) are strong indicators of the amount of chemical erosion from the continents. Unfortunately, these deep-sea records are generally unavailable for the Paleozoic. Ronov et al. (1980) have analyzed sedimentation patterns from shallow-water environments throughout the Phanero- zoic, but their two-part division of Early and Late Permian is too broad to isolate the increasing and decreasing portions of the curve. They found almost no change in average sediment accumulation rates between these divisions of the Permian; however, they did find a substantial decrease in the percent carbonate deposited (by volume), accompanied by a large increase in the percent marine elastics, in the Late Per- mian. For shallow-water environments this change in lithologies is consistent with increased weathering and erosion from the continents.

Thus we believe that the evidence for changes in climate, and hence weathering rate, during the Permian supports the conclusion that changes in the riverine Sr flux as well as the isotopic ratio of that flux are respon- sible for changes in oceanic *‘Sr/%r during this time. There is no need to invoke major changes in ocean circulation as has been suggested by other authors. For example, Gruszczynski et al. ( 1992) argued that under stagnant, stratified conditions the surface waters of the ocean would become increasingly enriched in radio- genie Sr from river runoff, while the deep waters received less radiogenic hydrothermal fluids. Sudden mixing of these water masses would rapidly reduce the 87Sr/86Sr value of the surface waters. In this scenario, the isotopic record preserved from shallow,water envi-

E. E. Martin, J.D. Macdougall / Chemical Geology 125 (1995) 73-99 95

ronments, such as those sampled for this study, would initially record a gradual increase followed by a sudden decrease, unlike the pattern illustrated by our Sr isotope data. Holser (1977) suggested a scenario in which rapid ocean mixing of an isolated brine pool could explain the rapid rise of seawater S34S in the Early Triassic; however, the Sr isotope minimum preceded the S34S minimum by several million years and the two events are clearly separated by the P/Tr boundary.

6. Conclusions

A detailed record of seawater 87Sr/86Sr for the Per- mian indicates that a minimum value for Phanerozoic time occurred in the Capitanian stage of the Late Per- mian. The isotopic composition of seawater for 10 Ma prior to this point decreased at an average rate of 0.000062 Ma-‘, while values from the Capitanian into the Triassic increased at an average rate of 0.000097 Ma- ‘. This increase is roughly equivalent to the steep- est portions of the seawater 87Sr/86Sr curve for the past 40 Ma, and is approximately two and a half times greater than the aver,age increase over that time period.

Model results indicate that these variations were pri- marily caused by fluctuations in the riverine flux of Sr during the decreasing segment of the curve, and a com- bination of a change in the riverine flux and the isotopic ratio of that flux to the oceans during the increasing segment. Limited Nd isotope data support the theory that the isotopic ratio of the flux became more “conti- nental” during the Late Permian; however, it is difficult to quantify the extent of this isotopic change. Conti- nental weathering raites probably increased in the Late Permian in conjunction with an increase in atmospheric CO*. Thus, in contrast to the large-scale seawater “Sr/ *%r increase in the Cenozoic, which appears to be controlled by weathering and climate changes associ- ated with uplift of ,the Himalaya, the increase in the Permian seems to have been influenced by changes in weathering associatjed with other climate forcing fac- tors, such as an increase in the concentration of atmo- spheric COZ, which resulted in global warming and increased humidity. Erosion of mountains in the equa- torial region, and their affect on the development of monsoonal circulati,on and the transport of rain into the continental interiors,, may also have played a role.

Acknowledgements

We particularly want to thank B.R. Wardlaw (USGS, Reston, Virginia) for allowing us to dissolve precious conodont samples from impoverished P/Tr fauna. He generously provided all of the conodonts analyzed for this study, as well as advice on P/Tr stra- tigraphy. D. Clark kindly supplied several conodont specimens that were used for the leaching experiment. We also want to thank C. MacIssac for assistance in the laboratory, and G.W. Lugmair for analytical advice. D.H. Erwin, W.T. Holser and G. Retallack read and provided helpful comments on an early draft of this paper, and L. Derry plus two anonymous reviewers greatly improved the final version. This research was supported by NSF (EAR9 1 - 183 19).

References

Aldridge, R.J., 1987. Conodont palaeobiology: A historical review. In: R.J. Aldridge (Editor), Palaeobiology of Conodonts. Wiley, New York, N.Y., pp. 1 l-34.

Armstrong, R.L., 1971. Glacial erosion and the variable isotopic composition of strontium in sea water. Nature (London), 230: 132-133.

Baksi, A.K. and Farrar, E., 1991. 40Ar/39Ar dating of the Siberian Traps, USSR: Evaluation of the ages of the two major extinction events relative to episodes of flood-basalt volcanism in the USSR and the Deccan Traps, India. Geology, 19: 4614.

Barnes, C.R., Sass, D.B., and Monroe, E.A., 1973. Ultrastructure of some Ordovician conodonts. Geol. Sot. Am. Spec. Pap., 141: l- 30.

Baud, A., Magaritz, M. and Holser, W.T., 1989. Permian-Triassic of the Tethys: Carbon isotope studies. Geol. Rundsch., 78: 649- 677.

Bemat, M., 1974. Les isotopes de l’uranium et du thorium et les terres rares dans l’environnement marine. Cah. OSTROM (Off. Sci. Tech. Rech. Outre-mer), S6. GBol., 7: 65-83.

Bemer, R.A., 1990. Atmospheric carbon dioxide levels over Pha- nerozoic time. Science, 249: 1382-1386.

Bemer, R.A., 1991. A model for atmospheric CO2 over Phanerozoic time. Am. J. Sci., 291: 339-376.

Bertram, C.J. and Eldefield, H., 1993. The geochemical balance of the ran? earth elements and neodymium isotopes in the oceans. Geochim. Cosmochim. Acta, 57: 1957-1986.

Bertram, C.J., Elderfield, H., Aldridge, R.J. and Morris, SC., 1992. “Sr/‘%r, “‘Nd/‘44Nd and REEs in Silurian phosphatic fossils. Earth. Planet. Sci. Lett., 113: 239-249.

Birck, J.L., 1986. Precision K-Rb-Sr isotopic analysis: Application to Rb-Sr chronology. Chem. Geol., 56: 73-83.

Brady, P.V. and Carroll, S.A., 1994. Direct effects of CO2 and tcm- pcrature on silicate weathering: Possible implications for climate control. Geochim. Cosmochim. Acta, 58: 1853-1856.

96 E.E. Martin, J.D. Macdougall / Chemical Geology 125 (I 995) 73-99

Briggs, D.G.,Clarkson,E.K.andAldridge,R.J., 1983,Theconodont animal. Lethia, 16: 1-14.

Burke, W.H., Denison, R.E., Hetherington, E.A., Koepnick, R.B., Nelson, H.F. and Otto, J.B., 1982. Variation of seawater 87Sr/ *%r throughout Phanerozoic time. Geology, 10: 516-519.

Caldeira, K., 1992. Enhanced Cenozoic chemical weathering and the subduction of pelagic carbonate. Nature (London), 357: 578- 581.

Campbell, I.H., Czamanske, G.K., Fedorenko, V.A., Hill, RI. and Stepanov, V., 1992. Synchronism of the Siberian Traps and the Permian-Triassic boundary. Science, 258: 1760-1763.

Capo, R.C. and DePaolo, D.J., 1990. Seawater strontium isotopic variation from 2.5 million years ago to the present. Science, 249: 51-55.

Caputo. M.V. and Crowell, J.C., 1985. Migration of the glacial cen- ters across Gondwana during the Paleozoic Era. Geol. Sot. Am. Bull., 96: 1020-1036.

Carr, T.R. and Paull, R.K., 1983. Early Triassic stratigraphy and paleogeography of the Cordilleran Miogeocline. In: M.W. Reyn- olds and E.D. Dolly (Editors), Mesozoic Paleogeography of the West-central United States. Rocky Mtn. Sect., Sot. Econ. Paleon- tol. Mineral., pp. 39-55.

Claoue-Long, J.C., Zichao, Z., Guogan, M. and Shaohua, D., 1991. The age of the Permian-Triassic boundary. Earth Planet. Sci. Lett., 105: 182-190.

Claypool, G.E., Holser, W.T., Kaplan, I.R., Sakai, H. and Zak. I., 1980. The age curves of sulfur and oxygen isotopes in marine sulfate and their mutual interpretation. Chem. Geol., 28: 199- 260.

Crowley, T.J., 1994. Pangean climates. In: G.D. Klein (Editor), Pangea: Paleoclimate, Tectonics and Sedimentation During Accretion, Zenith, and Breakup of a Supercontinent. Geol. Sot. Am. Spec. Pap., 288: 25-39.

Crowley, T.J. and North, G.R., 1991. Paleoclimatology. Oxford Uni- versity Press, Oxford, 339 pp.

Crowley, T.J., Hyde, W.T. and Short, D.A., 1989. Seasonal cycle variations on the supercontinent of Pangaea. Geology, 17: 457- 460.

Cummins, D.I. and Elderfield, H., 1994. The strontium isotopic com- position of Brigantain (late Dinantian) sea water. Chem. Geol., 118: 255-270.

Dahymple, G.B., Czamanske, G.K., Lanphere, M.A., Stephanov, V. and Federenko, V., 1991. 40Ar/39Ar ages of samples from the Noril’sk-Talnakh ore-bearing intrusions and the Siberian flood basalts, Siberia. Eos (Trans. Am. Geophys. Union), 72: 570 (abstract).

Delaney, M.L., Popp, B.A., Lepzeher, C.G. and Anderson, T.F., 1989. Lithium-to-calcium ratios in modem, Cenozoic, and Pale- ozoic articulate brachiopod shells. Paleoceanography, 4: 681- 691.

Denison, R.E., Koepnick, R.B., Burke, W.H., Hetherington, E.A. and Fletcher, A., 1994. Construction of the Mississippian, Penn- sylvanian and Permian seawater s7Sr/s%r curve. Chem. Geol. (Isot. Geosci. Sect.), 112: 145-167.

DePaolo, D.J. and Ingram, B.L., 1985. High-resolution stratigraphy with strontium isotopes. Science, 227: 938-941.

Edmond, J.M., 1992. Himalayan tectonics, weathering processes, and the strontium isotope record in marine limestones. Science, 258: 1594-1597.

Elderheld, H., 1986. Strontium isotope stratigraphy. Palaeogeogr., Palaeoclimatol., Palaeoecol., 57: 71-90.

Elderlield, H. and Gieskes, J.M., 1982. Sr isotopes in interstitial waters of marine sediments from Deep Sea Drilling Project cores. Nature (London), 300: 493-497.

Elderfield, H. and Greaves, M.J., 1982. The ram earth elements in seawater. Nature (London), 296: 214-219.

Elderfield, H. and Pagett, R., 1986. Rare earth elements in ichthyo- liths: Variations with redox conditions and depositional environ- ment. Sci. Total Environ., 49: 175-197.

Elderheld, H., Hawkeswotth, C.J., Greaves, M.J. and Calve& S.E., 198 1. Rare earth element geochemistry of oceanic ferromanga- nese nodules and associated sediments. Geochim. Cosmochim. Acta, 45: 5 13-528.

Epshteyn, O.G., 1981. Late Permianice-marinedeposits ofthe Atkan Formation in the Kolyma River headwaters region, U.S.S.R.. In: M.J. Hambey and W.B. Harland (Editors), Earth’s Pre-Pleisto- cene Glacial Record. Cambridge University Press, Cambridge, pp. 27&273.

Erwin, D.H., 1993. The Greatest Paleozoic Crisis. Columbia Uni- versity Press, New York, N.Y., 327 pp.

Fischer, A.G., 1964. Brackish oceans and the cause of the Permo- Triassic marine faunaI crisis. In: A.E.M. Naim (Editor), Prob- lems in Paleoclimatology. Wiley, New York, N.Y., pp. 566-579.

Fomey, G.G., 1975. Permo-Triassic sea-level change. J. Geol., 83: 773-779.

Forster, S.C. and Warrington, G., 1985. Geochronology of the Car- boniferous, Permian and Triassic. In: N.J. SneIIing (Editor), The Chronology of the Geological Record. Geol. Sot. London Mem., 10: 99-l 13.

F&es, L.A., 1979. Climates Throughout Geologic Time. Elsevier, Amsterdam, 3 10 pp.

Gallegos, D.M. and Wardlaw, B.R., 1992. Asselian (Early Permian) conodont fauna from the Garden Valley Formation, Sulphur Springs Range, Eureka County, Nevada. Geol. Sot. Am., Rocky Mtn. Sect., Abstr., 24: 13.

Goldstein, S.J. and Jacobsen, S.B., 1987. The Nd and Sr isotopic systematics of river-water dissolved material: Implications for the sources of Nd and Sr in seawater. Chem. Geol. (Isot. Geosci. Sect.), 66: 245-272.

Goldstein, S.L. and O’Nions, R.K., 1981. Nd and Sr isotopic rela- tionships in pelagic clays and ferromanganese deposits. Nature (London), 292: 324-327.

Gordon, W.A., 1975. Distribution by latitude of Phanerozoic evap- orite deposits. J. Geol., 83: 671-684.

Grossman, E.L., 1994. The carbon and oxygen isotope record during the evolution of Pangea: Carboniferous to Triassic. In: G.D. Klein (Editor), Pangea: Paleoclimate, Tectonics and Sedimentation During Accretion, Zenith, and Breakup of a Supercontinent. Geol. Sot. Am. Spec. Pap., 288: 207-228.

Gruszczynski, M., Halas, S., Hoffman, A. and Malkowski, K., 1989. A brachiopod calcite record of the oceanic carbon and oxygen isotope shifts at the Permian/Triassic transition. Nature (Lon- don), 337: 64-68.

E.E. Martin, J.D. MacdougaN / Chemical Geology 125 (1995) 73-99 97

Gruszczynski, M., Hoffman, A., Malkowski, K. and Veizer, J., 1992. Seawater strontium isotopic perturbation at the Permian-Triassic boundary, West Spitsbergen, and its implications for the inter- pretation of strontium isotopic data. Geology, 20: 779-782.

Hallam, A., 1977. Secular changes in marine inundation of USSR and North America through the Phanerozoic. Nature (London), 269: 770-772.

Hallam, A., 1984. Pre-Quatemary sea-level changes. Annu. Rev. Earth Planet. Sci., 12: 205-243.

Hallam, A., 1992. Phanerozoic Sealevel Changes. Columbia Uni- versity Press, New York, N.Y., 266 pp.

Harland, W.B., Armstrong, R.L., Cox, A.V., Craig, L.E., Smith, A.G., and Smith, D.G., 1990. A Geologic Time Scale 1989. Cambridge University Press, Cambridge, 263 pp.

Hays, J.D. and Pitman III, W.C., 1973. Lithospheric plate motion, sea-level changes an,d climatic and ecological consequences. Nature (London), 246: 18-22.

Hess, J., Bender, M. and Schilling, J.-G., 1986. Seawater “Sr/=Sr evolution from Cretateous to present. Science, 231: 979-984.

Hodell, D.A., Mueller, P. A., McKenzie, J.A. and Mead, G.A., 1989. Strontium isotope stratigraphy and geochemistry of the late Neo- gene. Earth Planet. Sci. L&t., 92: 165-178.

Hodell, D.A., Mead, G.A. and Mueller, P.A., 1990. Variation in the strontium isotopic composition of seawater (8 Ma to present) : Implications for chemical weathering rates and dissolved fluxes to the oceans. Chem. Geol. (Isot. Geosci. Sect.), 80: 291-307.

Holser, W.T., 1977. Catastrophic chemical events in the history of the ocean. Nature (London), 267: 403-408.

Holser, W.T. and MagaEritz, M., 1987. Events near the Permian- Triassic boundary. Mod. Geol., 11: 155-180.

Holser, W.T., Magaritz, M. and Clark, D.L., 1986. Carbon-isotope stratigraphic correlations in the Late Permian. Am. J. Sci., 286: 390-402.

Holser, W.T., Schanlaub, H.-P., Attrep, M., Jr., Boeckehnann, K., Klein, P., Magaritz, M., Orth, C.J., Fenninger, A., Jenny, C., Kralik, M., Mauri&h, H., Pak, E.. Schramm, J.-M., Stattegger, K. and Schmijller, R., 1989. A unique geochemical record at the Permian/Triassic boundary. Nature (London), 337: 39-44.

Holser, W.T., Schijnlaub, H.-P., Boeckelmann, K. and Magaritz, M., 199 1. The Permian-Triassic of the Gartnerkofel- I core (Carnic Alps, Austria): Synthesis and conclusions. Abh. Geol. B.-A., 45: 213-232.

Horita, J., Friedman, T.J., Lamar, B. and Holland, H.D.. 1991. The composition of Permian seawater. Geochim. Cosmochim. Acta, 55: 417432.

J&del, C. and Peng, T.-H., 1989. Isotopic compositions of neodym- ium in Atlantic surface waters may signal water flow from the Pacific and Indian oceans. Eos (Trans. Am. Geophys. Union), 70: 1132 (abstract).

Koepnick, R.B., Denison, R.E., Burke, W.H., Hetherington, E.A. and Dahl, D.A., 1990. Construction of the Triassic and Jurassic portion of the Phanerozoic curve of seawater “Sr/%r. Chem. Geol. (Isot. Geosci. Sect.), 80: 327-349.

Kominz, M., 1984. Oceanic ridge volumes and sea level change- an error analysis. Am. Assoc. Pet. Geol. Mem. 36: 109-127.

Kovach, J., 1980. Variations in the strontium isotopic composition of seawater during Phanerozoic time determined by analysis of conodonts. Geol. See. Am. Abstr., 12: 465 .

Kramm, U. and Wedepohl, K.H., 1991. The. isotopic composition of strontium and sulfur in seawater of Late Permian (Zechstein) age. Chem. Geol., 90: 253-262.

Krishnaswami, S., Trivedi, J.R., Satin, M.M., Ramesh, R. and Sharma, K.K., 1992. Strontium isotopes and rubidium in the Ganga-Brahmaputra River system; weathering in the Himalaya, fluxes to the Bay of Bengal and contributions to the evolution of oceanic 87Sr/86Sr. Earth Planet. Sci. L&t., 109: 243-253.

Kump, L.R., 1989. Alternative modeling approaches to the geochem- ical cycles of carbon, sulfur, and strontium isotopes. Am. J. Sci., 289: 390-410.

Kutzbach, J.E. and Gallimore, R.G., 1989. Pangaean climates: Mega- monsoons of the megacontinent. J. Geophys. Res., 94: 3341- 3357.

Magdtz, M., Bar, R.. Baud, A. and Holser, W.T., 1988. Thecarbon- isotopic shift at the Permian/Triassic boundary in the southern Alps is gradual. Nature (London), 331: 337-339.

Magaritz, M., Krishnamurthy, R.V. and Holser, W.T., 1992. Parallel trends in organic and inorganic carbon isotopes across the Per- mian/Triassic boundary. Am. J. Sci., 292: 727-739.

McLean, D.M., 1985. Deccan traps mantle degassing in the terminal Cretaceous marine extinctions. Cretaceous Res., 6: 235-259.

Menning, M., 1989. A synopsis of numerical time scales 1917-1986. Episodes, 12: 3-5.

Michard, A., Albtie, F., Michard, G., Minster, J.F. and Charlou, 1983. Ram-earth elements and uranium in high-temperature solu- tions from East Pacific Rise hydrothermal vent field ( 13”N). Nature (London), 303: 795-797.

Miller, K.G., Feigenson, M.D., Wright, J.D. and Clement, B.M., I99 1. Miocene isotope reference section, Deep Sea Drilling Pro- ject Site 608: An evaluation of isotopic and biostratigraphic res- olution. Paleoceanography, 6: 33-52.

Nalivkin,D.V., 1973.GeologyoftheU.S.S.R..UniversityofToronto Press, Toronto, Ont.. 855 pp.

Nie, S., Rowley, D.B. and Ziegler, A.M., 1990. Constraints on the locations of Asian microcontinents in Paleo-Tethys during the Late Permian. In: W.S. McKerrow and C.R. Scotese (Editors), Palaeozoic Palaeogeogmphy and Biogeography. Geol. Sot. Lon- don Mem., 12: 397-409.

Nishioka, S., Arakawa, Y. and Kobayashi, Y., 1991. Strontium iso- tope profile of Carboniferous-Permian Akiyoshi Limestone in southwest Japan. Geochem. J., 25: 137-146.

Odin, G.S., 1982. The Phanerozoic time scale revisited. Episodes, 1982: 3-9.

Odin, G.S., 1985. Comments on the geochronology of the Carbon- iferous to Triassic times. In: N.J. Snelling (Editor), The Chro- nology of the Geological Record. Geol. Sot. London Mem., 10: 114-116.

Palmer, A.R., 1983. The Decade of North American Geology: 1983 Geologic Time Scale. Geology, 11: 503-504.

Palmer, M.R. and Edmond, J.M., 1989. The strontium isotope budget of the modem ocean. Earth Planet. Sci. L&t., 92: 1 l-26.

Palme.r, M.R. and Edmond, J.M., 1992. Controls over the strontium isotope composition of river water. Geochim. Cosmochim. Acta, 56: 2099-211 I.

Palmer, M.R. and Elderfield, H., 1985. Sr isotope composition of sea water over the past 75 Myr. Nature (London), 314: 526-528.

98 E.E. Martin, J.D. Macdougall/Chemical Geology 125 (1995) 73-99

Parrish, J.T., 1993. Climate of the supercontinent Pangea. J. Geol., 101; 215-233.

Piepgras, D.J. and Wasserburg, G.J., 1985. Strontium and neodym- ium isotopes in hot springs on the East Pacific Rise and Guaymas Basin. Earth Planet. Sci. L&t., 72: 341-356.

Piepgras, D.J. and Wasserburg, G.J., 1987. Rare earthelement trans- port in the western North Atlantic inferred from Nd isotopic observations. Geochim. Cosmochim. Acta, 51: 1257-1271.

Piepgras, D.J., Wasserburg, G.J. and Dasch, E.J., 1979. The isotopic composition of Nd in different ocean masses. Earth Planet. Sci. Len., 45: 223-236.

Pietzner, H., Vahl, J., Werner, H. and Ziegler, W., 1968. Zur chem- ischen Zusammensetzung und Mikromorphologie der Conodon- ten. Palaeontographica, 128: 115-152.

PJRG (Pakistani-Japanese Research Group), 1985. Permian and Triassic systems in the Salt Range and Surghar Range, Pakistan. In: K. Nakazawa and J.M. Dickins (Editors), The Tethys -Her Paleogeography and Paleobiogeography from the Paleozoic to Mesozoic. Tokai University Press, Tokyo, pp. 219-312.

Popp, B.N., Anderson, T.F. and Sandberg, P.A., 1986a. Brachiopods as indicators of original isotopic compositions in some Paleozoic limestones. Geol. Sot. Am. Bull., 97: 1262-1269.

Popp, B.N., Podosek, F.A., Brannon, J.C., Anderson, T.F. and Pier, J., 1986b. s’Sr/%r ratios in Permo-Carboniferous sea water from the analyses of well-preserved brachiopod shells. Geochim. Cosmochim. Acta, 50: 1321-1328.

Raup, D.M. and Sepkoski, J.J., 1986. Periodic extinctions of families and genera. Science, 231: 833-836.

Raymo, M.E. and Ruddiman, W.F., 1992. Tectonic forcing of late Cenozoic climate. Nature (London), 359; 117-122.

Raymo, M.E., Ruddiman. W.F. and Froelich, P.N., 1988. Influence of late Cenozoic mountain building on ocean geochemical cycles. Geology, 16: 649-653.

Renne, P.R. and Basu, A.R., 1991. Rapid eruption of the Siberian Traps flood basalts at the Permo-Triassic boundary. Science, 253: 176-179.

Richter, F.M., Rowley, D.B. and DePaolo, D.J., 1992. Sr isotope evolution of seawater: the role of tectonics. Earth Planet. Sci. L&t., 109: 1 l-23.

Robinson, P.L., 1973. Palaeoclimatology and continental drift. In: D.H. Tarling and S.K. Runcom (Editors), Implications of Con- tinental Drift to Earth Sciences, Vol. 1. Academic Press, New York, N.Y., pp. 451476.

Ronov, A.B., Khain, V.E., Balukhovshy, A.N. and Seslavinsky, K.B., 1980. Quantitative analysis of Phanerozoic sedimentation. Sediment. Geol., 25: 31 l-325.

Ross, C.A. and Ross, J.R.P., 1987. Late Paleozoic sea levels and depositional sequences. In: CA. Ross and D. Haman (Editors), Timing and Depositional History of Eustatic Sequences: Con- straints on Seismic Stratigraphy. Cushman Found. Foraminiferal Res., Washington, D.C., Spec. Publ. No. 24, pp. 137-149.

Ruppel, S.C., James, E.W. and Banick, J.E., 1993. High-resolution strontium isotope. chemostratigmphy of the Paleozoic using con- odonts. Geol. Sot. Am. Abstr., 25: A473.

Schopf, T.J.M., 1974. Permo-Triassic extinctions: Relation to sea- floor spreading. J. Geol., 82: 129-143.

Scotese, CR. and McKerrow, W.S., 1990. Revised world maps and introduction. In: W.S. McKerrow and C.R. Scotese (Editors), Palaeozoic Palaeogeography and Biogeography. Geol. Sot. Lon- don Mem., 12: l-21.

Sepkoski, Jr., J.J., 1982. Massextinctions in the Phanerozoicoceans: A review. Geol. Sot. Am. Spec. Pap., 190: 283-289.

Shaw, H.F. and Wasserburg, G.J., 1985. Sm-Nd in marine carbon- ates and phosphates: Implications for Nd isotopes in seawater and crustal ages. Geochim. Cosmochim. Acta, 49: 503-518.

Staudigel, H., Doyle, P. and Zindler, A., 1985. Sr and Nd isotope systematics in fish teeth. Earth Planet. Sci. Lett., 76: 45-56.

Stevens, C.H., 1977. Was development of brackish oceans a factor in Permian extinctions? Geol. Sot. Am. Bull., 88: 133-138.

Sweet, WC., 1992. A conodont-based high-resolution biostratigra- phy for the Permo-Triassic boundary interval. In: W.C. Sweet, Y. Zunyi, J.M. Dickins and Y. Hongfu (Editors), Permo-Triassic Events in the Eastern Tethys. Cambridge University Press, Cam- bridge, pp. 120-133.

Sweet, W.C., Zunyi, Y., Dickins, J.M. and Hongfu, Y., 1992. Permo- Triassic events in the eastern Tethys - an overview. In: W.C. Sweet, Y. Zunyi, J.M. Dickinsand Y. Hongfu, (Editors), Permo- Triassic Events in the Eastern Tethys. Cambridge University Press, Cambridge, pp. 1-8.

Taylor, E.L., Taylor, T.N. and Cdneo, N.R., 1992. The present is not the key to the past: A polar forest from the Permian of Antarctica. Science, 257: 1675-1677.

Teichett, C., 1990. The Permian-Triassic boundary revisited. In: E.G. Kauffman and O.H. Wall&r (Editors), Extinction Events in Earth History. Springer, New York, N.Y., pp. 199-238.

Vail, P.R., Mitchum, R.M., Todd, R.G., Widmier, J.M., Thompson, S., Sangree, J.B., Bubb, J.N. and Hatelid, W.G., 1977. Seismic stratigraphy and global change. Am. Assoc. Pet. Geol. Mem., 26: 49-212.

Veizer, J. and Compston, W., 1974. 87Sr/86Sr composition of sea- water during the Phanerozoic. Geochim. Cosmochim. Acta, 38: 1461-1484.

Wardlaw, B.R., 1995. Permian conodonts. In: P.A. Scholle and P.N. Peryt (Editors), Permian of Northern Pangaea, Vol. 1: Paleo- geography, Paleoclimates and Stratigraphy. Springer, New York, N.Y., pp. 186195.

Wardlaw, B.R. and Collinson, J.W., 1986. Paleontology and depo- sition of the PhosphoriaFormation. Contrib. Geol., 24: 107-142.

Wardlaw, B.R. and Grant, R.E., 1990. Conodont biostratigraphy of the Permian Road Canyon formation, Glass Mountains, Texas. U.S. Geol. Surv. Bull., 1895: Al-A18.

Wardlaw, B.R. and Pogue, K.R., 1995. The Permianof Pakistan. In: P.A. Scholle and P.N. Peryt, (Editors), Permian of Northern Pangaea, Vol. 2: Sedimentary Basins and Economic Resources. Springer, New York, N.Y., pp. 23-40.

Wardlaw, B.R., Davis, R.A., Rohr, D.M., and Grant, R.E., 1990. Leonardian-Wordian (Permian) deposition in the northern Del Norte Mountains, West Texas. U.S. Geol. SW. Bull., 1881: Al- A14.

Wardlaw, B.R., Grant, R.E. and Rohr, D.M., 1991. Proc. Guadalu- pian Symp., Alpine, Texas, 111 pp.

Wignall, P.B.andHallam, A., 1993.Griesbachian (EarliestTriassic) palaeoenvironmental changes in the Salt Range, Pakistan and

E.E. Martin, J.D. Macdougall / Chemical Geology 125 (1995) 73-99 99

southeastern China and their bearing on the Permo-Triassic mass extinction. Palaeogeogr., Palaeoclimatol., Palaeoecol., 102: 215- 231.

Wright, J., Seymour, R.S. and Shaw, H.F., 1984. REE and Nd iso- topes in conodont apatite: Variations with geological age and depositional environment. In: D.L. Clark (Editor), Conodont Biofacies and Provinc:ialism. Geol. Sot. Am. Spec. Pap., 196: 325-340.

Nachos, J.C., 1993. Onset of southern hemisphere glaciation in the Early Cenozoic and the Sr isotopic composition of seawater. Eos (Trans. Am. Geophys. Union), 74: 177 (abstract).

Zharkov, M.A., 1981. History of Paleozoic Salt Accumulation. Springer, New York, N.Y., 308 pp.

Zhou, L. and Kyte, F.T., 1988. The Permian-Triassic boundary event: A geochemicaJ study of three Chinese sections. Earth Planet. Sci. Lett., 90: 411-421.

Wright, J., Schrader, H. and Holser, W.T., 1987. Paleoredox varia- Ziegler, P.A., 1989. Evolution of Laumssia: A Study in Late Paleo- tions in ancient oceans recorded by rare earth elements in fossil zoic Plate Tectonics. Kluwer, Dordrecht. apatite. Geochim. Cosmochim. Acta, 51: 631644. Zonenshain, L.P., Korinevsky, V.G., Kazmin, V.G., Pechersky,

Yemane, K., 1993. Contribution of Late Permian paleogeography in D.M., Khain, V.V. and Matveenkov, V.V., 1984. Plate tectonic maintaining a temperate climate in Gondwana. Nature (Lon- model of the south Urals development. Tectonophysics, 109: 95- don), 361: 51-54. 135.