Palaeogeography, Palaeoclimatology, Palaeoecology 285 (2010) 307–320
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Palaeogeography, Palaeoclimatology, Palaeoecology
j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / p a l a e o
Carbon and oxygen isotope evidence for high-frequency (104–105 yr) and My-scale
glacio-eustasy in Middle Pennsylvanian cyclic carbonates (Gray Mesa Formation),
central New Mexico
Maya Elrick ⁎, Lea Anne Scott
Earth and Planetary Sciences, University of New Mexico, Albuquerque NM 87131, United States
a r t i c l e
i n f o
Article history:
Received 16 July 2009
Received in revised form 5 November 2009
Accepted 17 November 2009
Available online 22 November 2009
Keywords:
Pennsylvanian
Carbon and oxygen isotopes
Glacio-eustasy
Conodonts
Icehouse climates
a b s t r a c t
We combine cyclo- and sequence stratigraphy along with whole rock δ13C and conodont apatite δ18O
analysis to document high-frequency (104–105 yr) and My-scale sea-level changes for the Middle
Pennsylvanian (Desmoinesian or Moscovian) Gray Mesa Formation of central New Mexico. Approximately
75 subtidal cycles (1–8 m) are grouped into 4 1/2 My-scale depositional sequences (40–80 m). About 50% of
the cycles show evidence of prolonged subaerial exposure at cycle tops with the development of calcretes,
diagenetic mottling, and regolith intraclasts. High-resolution δ13C analysis of whole rock limestones across
nine of the cycles indicates that the cycle tops were diagenetically altered by isotopically light, meteoric
fluids during sea-level fall and lowstand. These δ13C trends support the interpretation that high-frequency
sea-level changes were responsible for cycle development.
Conodont apatite δ18O values from sampled cycles indicate that the high-frequency sea-level changes were
driven by glacio-eustasy combined with changes in surface seawater temperature (SST). δ18O values from
conodont apatite, spanning parts of three depositional sequences indicate that My-scale glacio-eustasy and/or
SST changes controlled sequence development. δ18O shifts indicate that the magnitudes of 104–105 yr glacioeustasy were between ∼ 55 and 170+ m combined with tropical SST changes of ∼ 1.5°–6 °C. Calculated My-scale
glacio-eustatic oscillations were between ∼ 60 and 140 m with SST changes of b 3.5 °C. The most plausible
driver for the My-scale paleoclimate changes is long-period obliquity (∼ 1.2 My) variations. These calculated
high-frequency, glacio-eustatic values are similar or greater than Pleistocene values, and lie within the range
estimated for other Middle Pennsylvanian successions using a variety of independent eustatic proxies. The
similarity in range of magnitudes between high-frequency and My-scale sea-level changes combined with the
large differences in magnitudes between individual high-frequency sea-level oscillations helps explain the
lack of systematic cycle-stacking patterns within these Pennsylvanian icehouse sequences.
© 2009 Elsevier B.V. All rights reserved.
1. Introduction
The Late Paleozoic (Carboniferous to Permian) ice age is well
documented by the occurrence of glacial deposits on all continental
remnants of southern Gondwana (Caputo and Crowell, 1985; Veevers
and Powell, 1987; Frakes and Francis, 1988; Crowell, 1999; Isbell et al.,
2003; Fielding et al., 2008). Traditional reconstructions of this glacial
period invoke a single, large, and long-lived (∼50–90 My) ice sheet (e.g.,
Veevers and Powell, 1987). More recently, the concept of discrete glacial
and nonglacial intervals each lasting on the order of several million
years and involving multiple ice sheets is unfolding (Isbell et al., 2003;
Montañez et al., 2007; Fielding et al., 2008).
Far-field (low-latitude) evidence of the waxing and waning of
these ice sheets comes from the globally widespread occurrence of
⁎ Corresponding author.
E-mail addresses: [email protected] (M. Elrick), [email protected]
(L.A. Scott).
0031-0182/$ – see front matter © 2009 Elsevier B.V. All rights reserved.
doi:10.1016/j.palaeo.2009.11.023
high-frequency (104–105 yr) cycles/parasequences (or cyclothems).
In the Pennsylvanian of the U.S. Midcontinent and United Kingdom,
these high-frequency cycles are typically composed of a combination
of offshore siliciclastics, nearshore skeletal limestones, fluvial-deltaic
siliciclastics, coals, and paleosols (e.g., Moore, 1964; Ramsbottom,
1979; Maynard and Leeder, 1992; Heckel, 1977, 1994; Wright and
Vanstone, 2001). In contrast, many Pennsylvanian marine successions in the U.S. Southwest are composed predominantly of offshore
through nearshore carbonates; nonmarine siliciclastics, paleosols,
subaerial exposure features can be either poorly exposed or relatively
weakly developed (Goldhammer and Elmore, 1984; Algeo et al.,
1991; Soreghan, 1994; Wiberg and Smith, 1994; Scott and Elrick,
2004).
The magnitude of sea-level changes associated with the development of these Pennsylvanian cycles has been estimated using a variety
of independent mechanisms including ice-volume modeling (Crowley
and Baum, 1991), facies juxtaposition (Heckel, 1994, 1997), paleotopography (Soreghan and Giles, 1999), and δ18O isotopic shifts (Adlis
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et al., 1988; Joachimski et al., 2006). In this paper, we 1) present highresolution δ13C trends associated within Middle Pennsylvanian cycles
and cycle-capping early diagenetic features to verify high-frequency
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sea-level changes, 2) present conodont apatite δ O trends from cycles
and My-scale (3rd-order) depositional sequences to document glacioeustasy as the driver of 104–105 yr and My-scale sea-level changes,
and 3) estimate the magnitudes of the glacio-eustatic changes and
discuss the implications these magnitudes have for cycle-stacking
patterns in icehouse climate modes.
2. Geologic setting and stratigraphy
The early assemblage of Pangea in the Pennsylvanian resulted in the
development of the Ancestral Rocky Mountains and associated
intracratonic basins across the western U.S. (Kluth and Coney, 1981;
Ye et al., 1996). During the Pennsylvanian, New Mexico lay within 5–10°
of the paleoequator and was covered by shallow tropical seas
interrupted by uplifted mountain blocks composed predominately of
Precambrian crystalline rocks (Fig. 1a).
Fig. 1. (a) Paleogeographic map of New Mexico during the Middle Pennsylvanian; inset of the western U.S. shows the location of New Mexico. Darker gray areas represent uplifts related to
formation of the Ancestral Rocky Mountains. Mesa Sarca study area in black box. Other Middle Pennsylvanian outcrops discussed in text include: MA = Mesa Aparejo, SM = Sandia
Mountains, SA = Sacramento Mountains. (b) Chronostratigraphic diagram for study area and most of central New Mexico. Vertical lines represent a hiatus. Gray Mesa Formation shown in
shaded area. Biostratigraphic ages and nomenclature compiled from Kelley and Wood (1946) and Kues (2001).
M. Elrick, L.A. Scott / Palaeogeography, Palaeoclimatology, Palaeoecology 285 (2010) 307–320
The Lucero basin study area lay east of the low-relief Zuni uplift
and accumulated up to ∼ 300 m of Middle Pennsylvanian carbonates
and subordinate siliciclastics (Figs. 1a and 2). Northward thickening
and coarsening of the siliciclastics within this basin suggest that most
of them were sourced from the Penasco uplift lying north of
Albuquerque (Fig. 1a). Based on extrapolation of data presented by
Dickinson and Lawton (2003), subsidence rates in the Lucero basin
were highest during the Late Pennsylvanian (Virgilian or Kasimovian)
with tectonic movement occurring along high-angle normal faults
(Cather, 2001).
Within the study area, Pennsylvanian deposits lie unconformably
above Precambrian crystalline rocks and are composed of coarse- to
fine-grained transgressive siliciclastics (Sandia Formation), mixed
carbonate-siliciclastics (Madera Group), and are conformably overlain
by Permian fluvial red beds of the Bursum and Abo Formations
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(Fig. 1b). The Middle Pennsylvanian (Desmoinesian or Moscovian)
Gray Mesa Formation (∼ 290 m) is exceptionally well exposed at Mesa
Sarca in central New Mexico (Figs. 2 and 3). At this locality, individual
beds can be physically traced along depositional strike for over 4 km.
Age control on the Gray Mesa Formation is based on limited fusulinid
and preliminary conodont biostratigraphy (Kelley and Wood, 1946;
Wengerd, 1959; Martin, 1971, this study).
3. Cycles and sequences
3.1. Meter-scale cycles
Shallow subtidal through deep subtidal depositional facies are
recognized within the Gray Mesa Formation; detailed descriptions
and depositional interpretations of the various facies are reported in
Fig. 2. Generalized stratigraphic column of Gray Mesa Formation showing depositional sequences, boundaries of high-frequency cycles, and sea-level curve interpreted from
sequence stratigraphy. The sequence boundaries appear more abrupt that observed in the field because of scale of figure. Note the lack of thickening and thinning cycle-stacking
13
patterns and lack of systematic distribution of cycle-capping subaerial exposure features. See text for additional discussions. Filled triangles = cycles sampled for δ C analysis, open
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triangles = cycles sampled for δ O conodont apatite analysis.
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Fig. 3. Field photographs of Gray Mesa Formation at Mesa Sarca study site. (a) Depositional sequences 1–5 (white vertical lines) defined by slope-forming transgressive system tracts
(TST) and maximum flooding zones (MFZ) and cliff-forming highstand system tracts (HST). Entire outcrop is about 350 m thick. (b) Typical high-frequency cycle (#12) with thin
transgressive base (recessive under hang in shadow), thick-bedded shallow subtidal cap, and microkarst erosional top. Isotopic trends for this cycle shown in Figs. 5c and 7a. (c) Field
photograph of parts of Sequences 2 and 3 (thick white lines) and tops of internal high-frequency cycles (thin white lines).
Scott (2004) and Scott and Elrick (2004). The subtidal facies are
arranged into ∼ 75 m-scale (average thickness of ∼3 m) upwardshallowing cycles and form subtidal cycles (in contrast to peritidal
cycles which are capped by tidal flat or beach/foreshore facies).
Deeper subtidal cycles are characterized by poorly exposed calcareous
mudstone (deep subtidal facies) at the base overlain by skeletal
wackestones through packstones/rare grainstones (shallow subtidal
facies), whereas shallow subtidal cycles lack the calcareous mudstone
at their base and are composed entirely of skeletal wackestones
through packstones (Figs. 3–5). The cycles are asymmetric and record
Fig. 4. Field photographs of cycle-capping early diagenetic features. (a) Black laminated calcite or pedogenic calcrete (white arrows) infilling horizontal cracks within cycle cap.
(b) Reddish gray diagenetic mottles (white arrows) embedded in light gray, fine limestone matrix. (c) Subrounded to subangular intraclasts (white arrows) embedded in light gray,
fine limestone matrix.
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either entirely regressive deposition above the flooding surface or a
thin transgressive unit followed by a thick interval of regressive
deposits. Maximum flooding occurs near the base of cycles, though
often within poorly exposed calcareous mudstone intervals, so it is
not possible to observe the actual turnaround in depositional patterns.
Approximately 50% of the cycle tops display early diagenetic features
indicative of subaerial exposure (discussed below).
The average cycle duration is estimated by dividing the length of
the Desmoinesian/Moscovian by the number of observed cycles.
Depending on which time scale is used (Ross and Ross, 1988; Klein,
1990; Heckel, 2003; Gradstein et al., 2004), the average cycle duration
13
ranges from ∼50 to 100 ky (“high-frequency”). This duration is an
approximation because an unknown number of cycles occur within
the underlying, poorly exposed earliest Desmoinesian Sandia Formation, and the exact position of the Middle–Upper Pennsylvanian
boundary based on fusulinid biostratigraphy is poorly constrained in
the study area (Martin, 1971).
3.2. Early diagenetic features
Early diagenetic features at cycle caps and some cycle bases are
subtle, but pervasive in outcrop (Fig. 4) and include: 1) dark gray to
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Fig. 5. δ C trends from fine-grained matrix of high-frequency cycles. Refer to Fig. 2 for stratigraphic location of sampled cycles. (a) Cycles showing distinct up-cycle decreases in δ C
13
13
values. Some cycles show the lowest δ C values tens of cm below the cycle top. Note that the immediately overlying δ C value at base of cycle #39 shifts back to typical
13
Pennsylvanian marine values. (b) Cycles with no distinctive up-cycle trends in δ C values. Cycle #58 displays field evidence of subaerial exposure with diagenetic mottling and
13
intraclasts, and the δ C values of these cycle-capping diagenetic features are highly depleted relative to the immediately associated fine-grained matrix (see Fig. 6 and text for
additional discussion). The top of cycle #57 shows evidence of depletion throughout the sampled cycle top. (c) Three cycles, each sampled at two localities separated by ∼200 m,
13
show up-cycle variations in δ C trends and magnitudes indicating along-strike variations in early diagenesis. Cycles #17 and #13 show significantly different values along strike with
up-cycle depletions at one locality and not at the other. Tops of cycles #12 and #13 have black calcite, diagenetic mottling, intraclasts, or microkarstic erosion, but record up-cycle
13
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δ C depletion at only one locality. Note that all δ C values for the individual cycle-capping diagenetic features show depletion relative to the adjacent fine-grained host limestone.
M. Elrick, L.A. Scott / Palaeogeography, Palaeoclimatology, Palaeoecology 285 (2010) 307–320
313
Fig. 5 (continued).
black laminated calcite infilling vertical and horizontal cracks and
encrusting bed tops, 2) red- to orange-gray mottles within a gray to
yellow limestone matrix, and 3) subrounded to subangular limestone
intraclasts floating in a fine limestone matrix (Scott and Elrick, 2004).
Dark laminated calcite and mottling overprints primary depositional
textures and the intensity of their development decreases downwards
from the cycle tops, typically affecting only the upper few tens of cm
of the cycle. The intensity of diagenetic overprinting varies along
strike within a single layer, but the overprinting effects can be traced
at least 2 km along depositional strike. These field relationships
indicate that the cycle-capping features developed during early
diagenesis before deposition of the overlying beds and are related to
subaerial exposure. Half of the cycles (50%) display at least one cyclecapping early diagenetic feature, while nearly 20% of the cycles
display two or more features.
3.2.1. Black laminated calcite
Dark gray to black laminated calcite is observed in b20% of cycle
tops and occurs as discontinuous bands (mm to cm thick and up to
20 cm long), wispy stringers, patches, and bed–top encrustations
(Fig. 4a). The bands infill horizontal to vertical cracks, and the
stringers create an anastomosing network imparting a brecciated
appearance to the host limestone. Thick chert bands and cement-filled
fenestrae are also associated with the dark calcite. In thin section, the
calcite is characterized by dark brown laminated micrite which coats
grains, lines cavities, and is associated with alveolar–sepal structures,
dilation cracks, and rhizoliths.
3.2.2. Diagenetic mottling
Approximately 25% of the cycles are capped by diagenetic
mottling, which is characterized by rounded irregular patches of
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reddish, orange, to dark gray limestone (mottles) surrounded by light
gray to yellow-gray limestone host rock (Fig. 4b). Mottles are 0.25–
10 cm in diameter and are circular to elongate in cross section; they
are similar in size to burrows observed throughout the Gray Mesa
Formation but differ by their discoloration. The red (most common) to
orange discoloration is most intense along the periphery of individual
mottles, overall mottle discoloration decreases downward from cycle
tops, and the intensity of discoloration varies along depositional
strike. In thin section, mottles and host limestone boundaries are
diffuse with mottles composed of slightly more interparticle cement
and micron-size hematite crystals.
3.2.3. Intraclasts
Intraclasts occur at the tops of b20% of the cycles and are
characterized by subangular to subrounded dark gray limestone
clasts embedded in a light gray to yellow-gray limestone matrix
(Fig. 4c); both clast-supported and matrix-supported textures are
observed. Clasts are generally poorly sorted ranging in size from
b1 cm to 14 cm, are composed of fine-grained limestone, and have
clearly defined edges with the associated host limestone. The clasts
may be silicified or reddened, with more intense reddening occurring
around clast edges. Intraclast beds display sharp irregular to planar
contacts with underlying limestones or can have gradational contacts.
Overprinting by diagenetic mottling sometimes makes it difficult to
distinguish between diagentic mottles and discolored rounded
intraclasts. In thin section, clast boundaries are irregular and sharp.
3.3. Depositional sequences
The Gray Mesa cycles are bundled into 4 1/2 depositional
sequences which are ∼ 40–80 m thick (Figs. 2 and 3). Transgressive
system tract (TST) through maximum flooding zones (MFZ) are
characterized by slope-forming intervals of deep subtidal cycles,
whereas highstand system tracts (HST) are composed predominatly
of cliff-forming, carbonate-dominated shallow subtidal cycles (Figs. 2
and 3; Scott and Elrick, 2004). No single stratigraphic horizon
representing the sequence boundary is detected, rather a sequence
boundary zone (SBZ) or zone of accommodation minimum is
recognized by the turnaround from carbonate-dominated shallow
subtidal cycles to siliciclastic-based, deep subtidal cycles. Sequence
durations range from ∼0.7 to 1.4 My (3rd-order) depending on which
Middle Pennsylvanian time scale is used (Ross and Ross, 1988; Klein,
1990; Heckel, 2003; Gradstein et al., 2004).
Fig. 2 illustrates the cycle-stacking patterns within depositional
sequences. Of particular interest is that unlike typical greenhouse
climate cycle-stacking patterns (e.g., Goldhammer et al., 1993), the
Gray Mesa cycles do not display systematic thickening and thinning
patterns defining My-scale accommodation trends, nor are cycles
displaying cycle-capping exposure features concentrated near the
accommodation minima. Instead, the exposure features occur
throughout the sequences and some of the thinnest cycles occur
within TST and MFZ intervals.
The distinctive slope- versus cliff-forming outcrop patterns
defining sequences can be traced across the length of Mesa Sarca
(N4 km) and into the next outcrop belt approximately 15 km to the
north (Mesa Aparejo; Fig. 1a). In addition, coeval marine carbonates in
the Sandia Mountains N100 km to the northeast and Sacramento
Mountains ∼250 km to the south (Fig. 1a) are comprised of four to
five Middle Pennsylvanian depositional sequences (Algeo et al., 1991;
Wiberg and Smith, 1994; Krainer and Lucas, 2004). The similarity in
cyclic facies types, sequence number, and sequence thickness suggests
that Middle Pennsylvanian 3rd-order accommodation space trends
are regional in scale, rather than due to local variations in tectonically
driven subsidence. Ongoing sequence stratigraphic and biostratigraphic studies will aid in determining the specific regional extent of
sequence correlations.
4. Methods
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Samples for δ C analysis were collected from subtidal cycles and
cycle-capping diagenetic features to assess whether the subtidal
marine limestones were altered by meteoric fluids during sea-level
fall and lowstand. Nine of the 75 cycles were sampled; two of the
sampled cycles lack cycle-capping diagenetic features and were
sampled to evaluate whether subaerial exposure could be detected
13
from δ C isotopic trends alone. In addition, three of the cycles were
sampled at sites separated by N200 m to assess lateral variations in
13
δ C trends. Cycle-capping black laminated calcite, diagenetic mottles,
and intraclasts along with their immediately associated host limestone were sampled at 16 different stratigraphic horizons to evaluate
the cm-scale effects of subaerial exposure and diagenesis.
13
A total of 275 samples were analyzed for δ C analysis. Samples
were collected every 30–50 cm in the lower parts of cycles and every
5–20 cm in the upper parts. Fine-grained matrix (composed dominantly of pellets, microspar, and micritic/microsparitic cement) was
collected from thick section billets with a Dremel tool and 1 mm-wide
drill bit to avoid obvious skeletal fragments, coarse cements, and veins
and to obtain the necessary resolution between mottles, intraclasts,
and host limestone.
Oxygen isotopes from conodont apatite were sampled across two
cycles and across parts of three depositional sequences to evaluate the
origins of high-frequency and My-scale sea-level changes. Sequencescale samples were collected within the deepest water facies of ten
successive cycles to minimize the effects of high-frequency isotopic
variation related to cycle development.
For stable isotopes of carbonate samples, CO2 gas was extracted by
reaction with 100% phosphoric acid on a Thermoquest-Finnigan Gas
Bench II automated preparation device and was measured using a
Finnigan Delta Plus XL continuous flow isotope ratio mass spectrometer.
Analyses are reported relative to VPDB (Vienna Peedee belemnite) and
were routinely analyzed and compared to NBS 19 and Carrara Marble
13
standards. Average standard deviations on standards are ±0.1‰ for δ C
18
values and 0.2‰ for δ O values and reported relative to PDB.
Samples for conodonts were processed using standard conodont
concentration techniques (Sweet and Harris, 1989). Once the
conodonts were concentrated, the apatite was converted to Ag3PO4
using a modified version of O'Neil et al. (1994) to ensure that only
phosphorous-bound oxygen was analyzed. The Ag3PO4 crystals were
analyzed in a high temperature TC-EA reduction furnace at 1450 °C in
a He stream (e.g., Elrick et al., 2009). Isotope ratios of resultant CO are
measured in continuous flow isotope ratio mass spectrometry using a
Finnigan Mat 252 mass spectrometer. The precision of analyses based
on long-term measurement of standards is b0.3‰ and is monitored by
multiple analyses of several phosphate standards interspersed with
samples. Analyses are reported relative to SMOW.
5. Results
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5.1. Whole rock δ Ocarb
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Whole rock carbonate δ O values from the cycles range from −1.0
to −9.9‰ (average = − 6.0‰). The average values are significantly
18
lower than estimated δ O Pennsylvanian seawater values (−1.5 and
−3.0‰; Brand, 1982; Grossman et al., 1991; Algeo et al., 1992; Mii et
al., 1999) and show no systematic trends with upward-shallowing
13
facies patterns or with δ C trends. Therefore, we interpret that the
18
δ Ocarb values reflect the effects of diagenetic alteration during late
burial.
13
5.2. Whole rock δ C
Estimated Middle Pennsylvanian marine δ13C values are between
∼3‰ and 5‰ (Brand, 1989; Grossman et al., 1991; Algeo, 1996; Mii
M. Elrick, L.A. Scott / Palaeogeography, Palaeoclimatology, Palaeoecology 285 (2010) 307–320
et al., 1999). Whole rock carbonate δ13C values from the Gray Mesa
subtidal cycles range from 4.1‰ to −4.6‰ with the lowest values
occurring in the upper b1 m of cycle tops. Six of the nine sampled
cycles show up-cycle δ13C depletions of between 2‰ to 4‰ (Fig. 5). In
half of these, the smallest isotopic values occur a few tens of cm below
the cycle top; in the other half, the smallest values lie at the cycle top.
13
δ C values from the immediately overlying cycle base shift back to
typical Pennsylvanian marine values (Fig. 5a). The two sampled cycles
which lack field identified, cycle-capping subaerial exposure features
13
(cycles #1 and 58) show no systematic trends in δ C values (Fig. 5b).
Two of the three cycles sampled ∼200 m apart (cycles #13 and 17)
13
show significant along-strike differences in δ C isotopic trends
(Fig. 5c).
13
Fig. 6 illustrates the δ C values of cycle-capping diagenetic features
and their immediately adjacent host limestone. Values from black
laminated calcite are consistently lower than their immediately
adjacent limestone host by between 0.3‰ and 6‰. In half the
13
sampled diagenetic mottles, the mottles have δ C values up to 3‰
lower than the adjacent host limestone; the other half have values
13
similar to their host. Intraclasts consistently have δ C values between
0.4‰ and 5‰ lower than their associated matrix. Of interest is that
13
δ C values from diagenetic features from some cycle tops show
depletion even when the whole rock values do not (cycle #58;
Fig. 5b).
13
315
18
5.3. Conodont δ Oapatite
18
The δ Oapatite values of the two sampled cycles range from ∼19‰
to 22‰ and isotopic values increase by ∼1.0‰ to 3.9‰ from the base
18
to the top of the cycles (Fig. 7a). While δ Oapatite values show
18
systematic up-cycle increases, δ Ocarb values from corresponding
whole rock limestone show no systematic variations (Fig. 7a)
18
supporting the interpretation that δ Oapatite values reflect primary
18
marine seawater, while δ Ocarb values reflect the effects of diagenesis.
18
δ Oapatite values spanning parts of Sequences 1, 2, and 3 range
from ∼17‰ to 19.5‰ (Fig. 7b). The values decrease and reach a
minimum within the TST/MFZ and early HST, increase and peak
within the middle HST, systematically shift to lower values within the
upper HST, and reach the lowest values in the TST/MFZ and early HST
of the overlying Sequence 3. The magnitude of isotopic shift from the
top of the underlying Sequence 1 to the early HST of Sequence 2 is
∼1‰ and from the HST of Sequence 2 to the HST of Sequence 3 is
∼2.4‰ (Fig. 7b).
The magnitude of these cycle- and sequence-scale isotopic shifts
are minimum values because the study area was located along the
inner to middle shelf; therefore, sediments representing final stages of
sea-level fall, lowstand, and the initial rise did not accumulate at this
location. Given this, the full extent of the oxygen isotopic shifts is not
recorded at this location.
Fig. 6. δ C trends for cycle-capping early diagenetic features black laminated calcite (pedogenic calcrete), diagenetic mottles, and intraclasts. All sampled calcretes and intraclasts
13
13
have δ C values that are less than that of the immediately adjacent fine-grained host limestone. The δ C values for diagenetic mottles are significantly less than the adjacent host in
half the sampled horizons illustrating the cm-scale heterogeneity of pore-fluid flow. See text for additional discussion.
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subaerially exposed, but cycle top diagenesis occurred in a rockdominated pore-fluid system or that the meteoric cement abundance
was low or variably distributed (Goldstein, 1991).
6. Discussion
13
6.1. Whole rock δ C isotopic trends
13
We interpret that the marine to near marine δ C values at cycle
bases and the lighter isotopic values in the upper portion of cycles
represent the effects of early diagenetic alteration by isotopically light
meteoric fluids during high-frequency sea-level falls/lowstands and
subaerial exposure (e.g., Allan and Matthews, 1982; Beeunas and
Knauth, 1985; Algeo et al., 1991; Goldstein, 1991; Joachimiski, 1994;
Algeo, 1996; Immenhauser et al., 2002; Railsback et al., 2003; Theiling
et al., 2007). In this interpretation, during sea-level fall and lowstand,
marine sediments were progressively exposed to vadose and phreatic
zone meteoric waters enriched in isotopically light soil CO2, and
13
micritic cement with low δ C values precipitated from these fluids. It
is likely that the entire cycle thickness (and more) was exposed to
meteoric waters during this time, but the isotopic signature of
resultant pore fluids in the lower parts of the cycles was dominated by
13
the dissolution of marine carbonates whose δ C values were
significantly higher than meteoric fluids (i.e., pore fluids were rockdominated); therefore, any early diagenetic micritic cements at cycle
bases record marine to near marine isotopic values. In contrast, the
13
δ C values of pore fluids in the upper portion of cycles were
dominated by isotopically light meteoric fluids (water-dominated
pore fluids); therefore, the early diagenetic micritic cements in cycle
13
13
tops record low δ C values. We suggest that the increase in δ C values
in the upper few tens of cm of some of the cycle tops may be the result
of micritic cements precipitated into pores spaces after exposure and
13
during the subsequent transgression. As a result, the whole rock δ C
values at some cycle tops are greater than underlying samples,
reflecting a mix between the meteoric values developed during
subaerial exposure and more marine values developed during the
13
subsequent transgression. Similar peak depleted δ C values lying
some depth below cycle tops have been reported for Pleistocene
through Ordovician cyclic successions (Allan and Matthews, 1982;
Joachimiski, 1994; Driese et al., 1994; Algeo, 1996; Immenhauser
et al., 2002; Railsback et al., 2003).
13
Along-strike variations in up-cycle δ C trends (Fig. 5c) indicate
lateral heterogeneities in 1) pore-fluid isotopic composition, flow
paths and flow rates, 2) sediment porosity/permeability, 3) abundance, distribution, and isotopic composition of meteoric cement
phases, 4) vegetation cover and abundance, and/or 5) local topography. Centimeter-scale diagenetic heterogeneities are also implied
given the large differences in δ13C values between host limestone and
immediately adjacent diagenetic mottles (Fig. 6). Similar along-strike
13
and cm-scale variations in δ C values are reported by Goldstein
(1991) and Theiling et al. (2007).
13
Those cycles which record constant marine or near marine δ C
values from top to bottom (cycles #1, 12, and 58) may indicate that
those particular cycles 1) were not subaerially exposed, 2) the
diagenetically altered portion of the cycle was eroded during
exposure and/or subsequent transgression, 3) little primary porosity
was available for precipitation of meteoric cements, and/or 4) early
diagenesis occurred in a rock-dominated pore-fluid system, therefore
13
the micritic cements record marine to near marine δ C values. The
fact that some cycles display field evidence of subaerial exposure, but
13
lack δ C evidence of meteoric diagenesis (cycle #12) and some cycles
13
show up-cycle δ C decreases at one locality but not ∼ 200 m away
(cycle #17), suggests that the marine sediments were, in fact,
18
6.2. Cycle-capping diagenetic features
6.2.1. Black laminated calcites
The laminated black calcites are interpreted as soil carbonates
(laminar calcretes) precipitated during pedogenesis and root calcification (Harrison and Steinen, 1978; Goldstein, 1988; Wright, 1994).
13
This interpretation is supported by their consistently low δ C values
(soil/plant-derived CO2), association with rhizoliths, alveolar–sepal
structures, and dilation cracks, and their similarity to previously
reported calcretes (e.g., Harris and Nuna, 1975; Harrison and Steinen,
1978; Goldstein, 1988; Tucker and Wright, 1994).
6.2.2. Diagenetic mottling
We interpret that the diagenetic mottles represent backfilling of
burrows by slightly more porous/permeable sediment, followed by
early micritic cementation from meteoric fluids during sea-level fall
and lowstand (Horbury and Quing, 2004). The abrupt difference in
13
δ C values between some of the mottles and adjacent host limestone
(Fig. 6) reflects the variable primary porosity/permeability and flow
paths of pore fluids.
6.2.3. Intraclasts
We interpret that the isotopically depleted intraclasts at cycle tops
represent soil regolith clasts developed during subaerial exposure and
meteoric diagenesis (e.g., Goldstein, 1988; Tucker and Wright, 1994).
13
6.2.4. Summary of δ C trends
13
The combined results from δ C trends spanning individual cycles
and trends from individudal cycle-capping diagenetic features clearly
supports our interpretation that cycles developed in response to sealevel changes and that during sea-level fall, lowstand, and early rise,
marine sediments were subaerially exposed and altered by meteoric
fluids.
18
6.3. Conodont δ Oapatite trends
Analyses of oxygen isotopes from marine calcite and apatite are a
well-established paleoclimate tool for evaluating variations in
seawater temperatures and isotopic changes related to the growth
and melting of continental glaciers. In particular, Joachimski et al.
(2006) analyzed conodont apatite from Upper Pennsylvanian
cyclothems/cycles in the U.S. Midcontinent and reported systematic
18
up-cycle positive δ O isotopic shifts associated with upwardshallowing facies trends. They interpret that up to 1.7‰ isotopic
shifts are the result of the glacial ice-volume effect combined with
cooling surface seawater temperatures. Using the Pleistocene as an
icehouse analog, they interpreted Late Pennsylvanian interglacial–
glacial sea-level changes of over 120 m combined with subtropical
surface water temperature changes of ∼2–4 °C.
18
We interpret that the δ Oapatite isotopic shifts associated with Gray
Mesa cycle and sequence development are also the result of glacioeustasy (ice-volume effect) and tropical seawater temperature
changes. Isotopic changes due to the effects of evaporation and/or
freshwater influx (“salinity effect”) are ruled out because the amount
of evaporation/freshwater influx required to explain the measured
18
Fig. 7. (a) δ Oapatite (from conodont apatite) trends for sampled high-frequency cycles. Invariant δ O values from whole rock carbonate are also shown for comparison in cycle #12,
18
18
which supports the interpreted primary seawater δ O values for conodont apatite. (b) δ Oapatite trends across parts of depositional Sequences 1, 2, and 3. The relative sea-level curve
is derived from sequence stratigraphic interpretations. TST = transgressive systems tract, MFZ = maximum flooding zone, HST = highstand systems tract, and SBZ = sequence
boundary zone. Note that the lowest isotopic values do not coincide with the interpreted deepest water facies (MFZ) of Sequences 2 or 3, rather they occur within the shallowing
18
phase or HST, suggesting that the δ O-derived sea-level curve provides a more accurate indication of eustasy than sequence stratigraphic interpretations. See text for additional
discussion. Symbols shown in Figs. 2 and 5.
M. Elrick, L.A. Scott / Palaeogeography, Palaeoclimatology, Palaeoecology 285 (2010) 307–320
317
318
M. Elrick, L.A. Scott / Palaeogeography, Palaeoclimatology, Palaeoecology 285 (2010) 307–320
isotopic variations would preclude the occurrence of open-marine
biota observed within these deposits. If the isotopic shifts were due to
seawater temperature changes alone, then the amount of thermal
expansion/contraction (assuming temperature-density relations of
seawater to be 1.9 × 10−2 vol.%/°C) would account for b10–15 m of
sea-level change, which is not sufficient to explain the observed facies
juxtapositions (i.e., deep subtidal facies overlain by subaerially
exposed shallow subtidal facies).
The ∼ 1‰ to 3.9‰ range of up-cycle isotopic shifts for the Gray
Mesa cycles are larger than that reported by Joachimski et al. (2006),
but are within the range observed in the glacial–interglacial
18
Pleistocene δ O record from planktonic foraminifera (∼0.5‰ to
3.5‰; Imbrie et al., 1984). In the Pleistocene, the smaller isotopic
shifts of between 0.5 and 1.0‰ shifts are related to precession-driven
(∼20 ky) climate cycles, while the larger (3.5+‰) shifts reflect
eccentricity-driven (∼100 ky) glacial–interglacial cycles. We suggest
18
that the wide range in δ O shifts between the two sampled
Pennsylvanian cycles also reflect the difference in paleoclimate
change between the various orbital frequencies.
If we use the Quaternary as an analog and assume that about 30% of
18
the Pennsylvanian δ O isotopic signal is the result of tropical surface
seawater temperature (SST) change and the remaining portion is due to
ice-volume effects (Fairbanks and Matthews, 1978; Fairbanks, 1989),
this suggests that the cycle-scale isotopic shifts of ∼1‰ to 3.9‰
represent tropical seawater temperature changes of ∼ 1.5°–6 °C,
respectively, and glacio-eustatic sea-level changes of ∼55–170+ m,
respectively. At the sequence-scale, the measured ∼1‰ to ∼2.4‰
variations suggest SST changes on the order of b3.5 °C and glacioeustatic changes between ∼60 and 140 m. In both cases, these are
minimum estimates because deposits representing time intervals of
maximum glaciation are not present at this inner to middle shelf
position.
Previous estimates of Middle Pennsylvanian glacio-eustatic
changes derived from several different independent proxies (ice18
volume modeling, facies juxtaposition, paleotopography, and δ O
isotopic shifts) range from ∼ 40 to 150 m (Adlis et al., 1988; Crowley
and Baum, 1991; Heckel, 1994, 1977; Soreghan and Giles, 1999;
Joachimski et al., 2006; Rygel et al., 2008). Our estimates based on
18
cycle-scale and sequence-scale δ O isotopic shifts lie within this
range, and though they are minimum estimates, they suggest that the
volume of Pennsylvanian glacial ice growth and melting may have
been larger than the Pleistocene.
Our estimated magnitudes of My sequence-scale glacio-eustatic
changes are similar to or slightly smaller than those calculated for
104–105 yr changes. If this is correct, and if our interpretation that the
magnitudes of sea-level change related to individual cycle development varies significantly, then this would explain why Pennsylvanian
cycle-stacking patterns are so different from typical greenhouse
stacking patterns.
High-frequency cycles bundled within many Cambrian, Ordovician,
Devonian, Jurassic, and Cretaceous greenhouse sequences are commonly characterized by thicker-than-average subtidal facies-dominated
cycles developed during My-scale transgressions and thinner-thanaverage, peritidal cycles (with increased occurrence of cycle-capping
exposure features) developed during My-scale regressions and lowstands (e.g., Oslger and Read, 1991; Crevello, 1991; Goldhammer et al.,
1993; Elrick, 1995; Grötsch, 1996; Lehmann et al., 1998; Lehrmann and
Goldhammer, 1999). Traditionally, these greenhouse stacking patterns
have been attributed to relatively small magnitude, high-frequency sealevel changes superimposed on larger magnitude, My-scale sea-level
changes (Goldhammer et al., 1993).
In the case of Pennsylvanian icehouse cycles, the large and variable
magnitude, high-frequency sea-level changes that are similar to or
larger than the superimposed My-scale sea-level changes result in
nonsystematic cycle-stacking patterns. For example, in the Gray Mesa
Formation, some of the thinnest cycles with most intensely developed
cycle-capping subaerial exposure features occur within the TST and
MFZ, and some very thick cycles occur in the HST and near the
accommodation minimum (Fig. 2; Scott and Elrick, 2004). Similar
nonsystematic patterns of cycle thickness, intracycle facies patterns,
and cycle-capping exposure features are reported for Middle and
Upper Pennsylvanian cycles in the Sandia Mountains of central New
Mexico (Wiberg and Smith, 1994) Orogrande Basin and Sacramento
Mountains of southern New Mexico (Algeo et al., 1991; Soreghan,
1994). These characteristics suggest that cycle thicknesses and
internal facies were governed by successive high-frequency accommodation trends rather than My-scale trends and contrasts with
stacking patterns of typical greenhouse successions (Read et al., 1995;
Lehrmann and Goldhammer, 1999; Barnett et al., 2002).
18
Of particular interest is that the sequence-scale δ O isotopic shifts
do not systematically co-vary with the sequence stratigraphically
derived sea-level curve (Fig. 7b). Isotopic values in the HST of
sequence 2 begin to significantly decrease (implying melting of glacial
ice and/or SST warming) before the facies record deepening. The
18
highest δ O values occur within the middle HST (rather than in the
SBZ), and the lowest values occur within the early HST (rather than
the MFZ). These trends suggest that our ability to interpret waterdepth changes from facies alone is not sensitive enough to generate
18
accurate sea-level curves. The δ O isotopic trends on the other hand,
do record the environmental changes and have the potential to
provide more precise indicators of glacio-eustatic changes. Additional
18
δ O isotopic data from underlying and overlying sequences at this
locality and regionally coeval sections is necessary to fully understand
and document the relationships between facies-derived water-depth
curves and paleoclimatic and global sea-level trends.
6.4. My-scale climate drivers
18
If our δ O isotopic interpretations are correct, then it implies that a
My-scale Pennsylvanian climate driver operated in tandem with highfrequency (104–105 year) climate changes to control cycle and
sequence development. Over the past decade, ∼ 1–3 My-scale climatic
cyclicity has been recognized in the oxygen and carbon isotope,
magnetic susceptibility, foraminifera, and lithologic records of
Cenozoic, Mesozoic, and Paleozoic marine successions (e.g., Lourens
and Hilgen, 1997; Herbert, 1997; Shackleton et al., 1999; Matthews
and Frohlich, 2002; Wade and Pälike, 2004; Coxall et al., 2005;
Mitchell et al., 2008; Elrick et al., 2009). In these cases, the My-scale
paleoclimate records have been attributed to long-period modulations of obliquity (∼1.2 My) and eccentricity (∼ 2.4 My) (Laskar et al.,
1993; Laskar, 1999; Laskar et al., 2004). We suggest that the My-scale
paleoclimate changes in the Middle Pennsylvanian were driven by
long-period obliquity variations which lead to glacial ice growth and
melting in southern Gondwana, My-scale glacio-eustatic sea-level
oscillations with magnitudes of many tens to over a hundred meters,
and the development of globally widespread 3rd-order depositional
sequences.
7. Conclusions
1) Deep subtidal through shallow subtidal facies of the Middle
Pennsylvanian Gray Mesa Formation are grouped into ∼75 highfrequency (104–105 yr) upward-shallowing subtidal cycles. Over
50% of the cycles are capped by early diagenetic features including
pedogenic calcretes, regolith intraclasts, and diagenetic mottling.
The cycles are stacked into 4 1/2 My-scale (3rd-order) depositional sequences (40–80 m).
13
2) High-resolution δ C stratigraphy across nine of the cycles
indicates that the cycle tops were diagenetically altered by
isotopically light meteoric fluids during high-frequency sea-level
fall and lowstand. The base of the cycles retain marine or near
13
marine δ C values due to diagenesis in rock-dominated pore
M. Elrick, L.A. Scott / Palaeogeography, Palaeoclimatology, Palaeoecology 285 (2010) 307–320
13
fluids. Cycle-capping early diagenetic features preserve δ C
isotopic evidence of substantial meteoric diagenesis during
13
subaerial exposure. These combined δ C isotopic patterns demonstrate that sea-level changes related to cycle development can
be detected even within carbonate successions displaying subtle
facies changes and/or development of subaerial exposure features.
18
3) δ O values of conodont apatite from two sampled cycles support the
hypothesis that the high-frequency sea-level changes were driven
18
by glacio-eustasy combined with changes in SST. δ O values from
conodont apatite spanning parts of three depositional sequences
also indicates that My-scale (3rd-order) glacio-eustasy and SST
changes controlled sequence development. The difference between
lithologically interpreted sea-level curves and those determined
18
from δ O trends suggests that oxygen isotopes are a more sensitive
indicator of glacio-eustatic changes. Using the Pleistocene icehouse
as a modern analog, estimates on the magnitudes of glacio-eustasy
for the sampled cycles are between ∼55 and 170+ m combined
with tropical SST changes of ∼1.5°–6 °C. Estimated sequence-scale
glacio-eustatic oscillations range from ∼60 to 140 m with SST
changes of b3.5 °C. The most plausible driver for the My-scale
paleoclimate changes is long-period obliquity (∼1.2 My) variations.
4) The calculated high-frequency glacio-eustatic values are similar or
greater than Pleistocene values, and lie within the range estimated
for other Middle Pennsylvanian successions using a variety of
independent eustatic proxies. If these glacio-eustatic estimates are
correct, then the similarity in range of magnitudes between highfrequency and My-scale sea-level changes combined with the large
differences in magnitudes between individual high-frequency sealevel oscillations helps explain the lack of systematic cycle-stacking
patterns within these Pennsylvanian icehouse sequences.
Acknowledgements
Partial funding for this project was provided by New Mexico
Geological Society, GSA and AAPG Grants-in-Aid, and the UNM Earth
and Planetary Sciences Department (to LAS), and NSF #EAR-0518205
(ME). We thank Mike Mechenbier of the Four Daughters Ranch for
land access. Many thanks for field and lab assistance by James Ashby,
Ian Atwell, Sarah Caldwell, Linnah Neidel, Michael Emms, Tammy
Lara, Cheryl Townsend, Leilani Ringkvist, and Viorel Atudorei. The
manuscript benefited from reviews by Dan Lehrmann.
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