Marine biodiversification in response to evolving phytoplankton stoichiometry

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Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 277 – 291 www.elsevier.com/locate/palaeo

Marine biodiversification in response to evolving phytoplankton stoichiometry Ronald E. Martin a,⁎, Antonietta Quigg b , Victor Podkovyrov c a b

Department of Geological Sciences, College of Marine and Earth Studies, University of Delaware, Newark, DE 19716, USA Phytoplankton Dynamics Laboratory, Departments of Marine Biology and Oceanography, Texas A&M University at Galveston, Galveston, TX 77551, USA c Institute of Precambrian Geology and Geochronology, Russian Academy of Sciences (RAS), St. Petersburg, 199034, Russia Received 20 December 2006; received in revised form 18 July 2007; accepted 2 November 2007

Abstract Diversification of the marine biosphere is intimately linked to the evolution of the biogeochemical cycles of carbon, nutrients, and primary productivity. A meta-analysis of the ratio of carbon-to-phosphorus buried in sedimentary rocks during the past 3 billion years indicates that both food quantity and, critically, food quality increased through time as a result of the evolving stoichiometry (nutrient content) of eukaryotic phytoplankton. Evolving food quantity and quality was primarily a function of broad tectonic cycles that influenced not just carbon burial, but also nutrient availability and primary productivity. Increasing nutrient availability during the middle-to-Late Proterozoic culminated in the production of food (phytoplankton biomass and fresh dead organic matter) with C:P Redfield ratios sufficient to finally promote geologically-rapid biodiversification during the Proterozoic–Phanerozoic transition. This resulted in further, massive nutrient sequestration into biomass that triggered positive feedback via nutrient recycling (bioturbation, mesozooplankton grazing) on phytoplankton productivity. Increasing rates and depths of bioturbation through the Phanerozoic suggest that nutrient recycling continued to increase. Increasing bioturbation and nutrient cycling appear to have been necessary to sustain the primary productivity and “energetics” (biomass, metabolic rates, and physical activity such as predation) of the marine biosphere because of the geologically-slow input of macronutrients like phosphorus from land and the continued sequestration of nutrients into marine and terrestrial biomass. © 2007 Elsevier B.V. All rights reserved. Keywords: Autotroph; Biodiversity; Biosphere; Energetics; Eukaryote; Food; Macroevolution; Nutrients; Phytoplankton; Primary productivity; Recycling; Stoichiometry

1. Introduction The succession of fossil faunas through the Phanerozoic has been inferred to indicate that the “energetics” ⁎ Corresponding author. Tel.: +1 302 831 6755; fax: +1 302 831 4158. E-mail address: [email protected] (R.E. Martin). 0031-0182/$ - see front matter © 2007 Elsevier B.V. All rights reserved. doi:10.1016/j.palaeo.2007.11.003

of the marine biosphere (biomass, metabolic rates, and physical activity such as predation) generally increased through geologic time (Vermeij, 1987, 1995; Bambach, 1993, 1999; Martin, 1995, 1996, 1998a,b, 1999, 2001, 2002; Allmon and Ross, 2001; see also Bambach et al., 2002; Aberhan et al., 2006; Kowalewski et al., 2006; Martin et al., 2006). For example, predators that are thought to have been metabolically active, such as

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Fig. 1. Environmental change and diversification of the marine biosphere. Adapted from sources cited in the text (Thayer, 1983; Martin, 1996, 2002, 2003; Katz et al., 2004, 2005; Martin et al., 2006). Acritarch diversity includes form taxa (Butterfield, 2004), but diversity patterns nevertheless appear robust (Katz et al., 2004, 2005). Positive carbon isotope ratios (δ13C) indicate enhanced marine and/or terrestrial photosynthesis, whereas negative ratios indicate decreased photosynthesis and/or input of isotopically-lighter 12C from various sources; snowflakes indicate approximate times of well-developed glaciers. Increasing strontium isotope (87Sr/86Sr) is used as a general qualitative indicator of increased continental weathering and runoff. Note parallel rise and decline of Paleozoic Fauna and green algal lineage, and diversification of Mesozoic and Modern faunas and red algal lineage, especially diatoms. Biomarkers (dinosteranes) presumably indicative of dinoflagellates are present during the early-tomiddle Paleozoic, but prasinophytes are thought to have dominated (Quigg et al., 2003; Falkowski et al., 2004; Katz et al., 2004, 2005); unequivocal fossil dinoflagellates finally appear in the Triassic, when dinosteranes also became prominent, suggesting a widespread evolutionary radiation of this group at this time (Fensome et al., 1999).

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ammonoids and, especially, jawed fishes, replaced nautiloid cephalopods and eurypterids beginning in the Late Devonian (Signor and Brett, 1984; Vermeij, 1987; Bambach, 1999; Fig. 1). Increasing energetics is thought to have been especially pronounced as increasingly modern marine faunas rose to prominence during the Meso-Cenozoic. These faunas are dominated by bivalves, gastropods, and teleost fish, modern representatives of which often have high metabolic rates or may be predatory. Modern representatives of the dominant extant taxa of each of these faunas also exhibit an exponential increase in weight-normalized (size-independent) respiration, supporting the hypothesis of an increase of metabolic rates through time (Zotin, 1984). The presumed increase of marine biosphere energetics has been attributed in part to increasing food availability (e.g., Bambach, 1993; Martin, 1996; Bambach, 1999). The trend of increasing food availability may be inferred from: 1) the increasing height of suspension feeding above the sediment–water interface by sessile epifauna during the Paleozoic, implying greater phytoplankton populations (food) above bottom (Bambach, 1993; Martin, 1996); 2) the later decline of the height of suspension feeding during the Mesozoic, most likely in response to increased cropping (Aberhan et al., 2006; Kowalewski et al., 2006; Madin et al., 2006; Wagner et al., 2006) by a more “energetic” benthos (Bambach, 1999); 3) the concomitant increase in the depth and intensity of burrowing (Thayer, 1983; Vermeij, 1987; Sepkoski et al., 1991); and 4) increasing hard part production by benthic invertebrates (Kidwell and Brenchley, 1994). Increasing food availability is also suggested by the succession of dominant phytoplankton taxa through time. The dominant phytoplankton taxa of the Proterozoic and Paleozoic were prasinophytes, chlorophytes,

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and possibly some early dinoflagellates, all assigned to the polyphyletic group “acritarchs” (Fig. 1). Dinoflagellates and coccolithophorid oozes dominated in the Mesozoic and Paleogene, followed by the explosive diversification of diatoms during the Neogene (Fig. 1; Martin, 1995, 1996; Katz et al., 2004, 2005). Modern representatives of each of these major taxa tend to prefer progressively more nutrient-rich conditions (Kilham and Kilham, 1980) and cultivated strains of each tend to be more phosphorus-rich (Table 1; Quigg et al., 2003), suggesting that food availability (carbon and nutrients) increased through the Phanerozoic (Martin, 1996, 2002). The evolution of phytoplankton stoichiometry in turn appears related to tectonic (Wilson) cycles of supercontinent rifting and reassembly and associated climate change (Falkowski et al., 2004; Katz et al., 2004, 2005). During the Phanerozoic, rifting continents are associated with broad sea-level rise, warm climate (high CO2), sluggish ocean circulation, and anoxia (Fig. 1). Acritarchs belong to “green” (chlorophyll-b bearing) phytoplankton lineages that use iron as a coenzyme in their photosynthetic pathways. Anoxia is therefore thought to have promoted the evolution of acritarchs during the early-to-middle Paleozoic by promoting the release of iron (and phosphorus) from bottom sediments (Falkowski et al., 2004; Katz et al., 2004, 2005). “Red” phytoplankton lineages (primarily coccolithophorids and dinoflagellates) expanded during the Mesozoic, with diatoms appearing in the Cretaceous, all presumably in response to increasing oxygenation (due to sequestration of carbon in marine sediments and land plants) and nutrient availability (Falkowski et al., 2004; Katz et al., 2004, 2005). Modern red phytoplankton taxa preferentially incorporate molybdenum (Mo), which is soluble in oxygenated water, into their photosynthetic apparatus (Table 1; Quigg et al., 2003; Falkowski et al.,

Table 1 Dominant phytoplankton taxa during geologic time and stoichiometric characteristics of modern representatives a Taxon

First appearance in geologic record

Modern C:P Ratio (based on cultivated strains)

Proterozoic (2.5–0.544 Ga): Prasinophyceae (Green) Chlorophyceae (Green)

1.2 Ga 1.0 Ga

∼ 200 ∼ 200

Paleozoic Era (544–0.245 Ga) Dinophyceae (Red)

0.44 Ga

∼ 140

0.21 Ga 0.12 Ga

∼ 50 ∼ 60

Meso-Cenozoic (0.245–0 Ga) Prymnesiophyceae (Red) (Coccolithophorids) Bacillariophyceae (Red) (Diatoms) a

From Quigg et al. (2003).

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2004; Katz et al., 2004, 2005). Diatoms became the primary exporters of organic carbon in the oceans when they diversified explosively in the Neogene in response to glaciation (which increased physical erosion and thus weathering rates), and the input of dissolved silica to the oceans that resulted from the spread of grasses (due to drying out of the Earth by glaciation) and sea-level fall (Quigg et al., 2003; Falkowski et al., 2004; Katz et al., 2004, 2005). Nevertheless, the wide ranges of δ13C (Fig. 1) and total organic carbon (TOC; Fig. 2A and B) have been inferred to indicate relatively constant primary productivity through geologic time (Schidlowski et al., 1983; Strauss et al., 1992). The role of evolutionary escalation as a driver of increasing global marine biodiversity has also been disputed (e.g., Madin et al., 2006). The apparent increase of marine biodiversity (Fig. 1) has been suggested to result from “the pull of the Recent” (i.e., the extension of the stratigraphic ranges of fossil taxa toward the Recent by more extensive sampling of Recent biotas; Raup, 1976, 1979). Differential preservation and sampling intensity have also been suggested to bias the fossil record of biodiversity toward the Recent (Alroy et al., 2001; Peters and Foote, 2001). However, recent studies suggest that the pull of the Recent has been overemphasized for Cenozoic bivalves (Jablonski et al., 2003). Also, Kidwell (2005) found no preservational bias of mollusc shell composition while Martin (2003) concluded that although the preservation of the sedimentary record is biased toward the present, this trend is confounded with concomitant trends of increasing nutrient levels, primary productivity, and food availability. So, did the energetics and food availability of the marine biosphere increase or not? Are there are other ways to assess the relationship of nutrient levels and primary productivity to food availability and energetics through geologic time? And what are the potential implications for the biodiversification and complexification of marine food webs and their relationship to the biogeochemical cycles of carbon and phosphorus? 2. Approach and rationale We examined primary productivity, carbon burial, and phosphorus availability through time by conducting

a meta-analysis of TOC (n = 2783) and phosphorus (n = 976) of sedimentary rocks of the past 3 billion years compiled from different sources. We also calculated the corresponding ratios of TOC buried to phosphorus (n = 885), here termed C:P burial ratios to distinguish them from Redfield C:P ratios of fresh organic matter. The rationale for this approach is as follows. Among the macronutrients necessary for biological growth is phosphorus, high concentrations of which occur in nucleic acids (DNA and RNA) that are critical to information storage and retrieval (Sterner and Elser, 2002). RNA, in particular, is necessary for the greatly increased rates of protein synthesis during rapid biomass growth and development (Sterner and Elser, 2002; Dell'Anno and Danovaro, 2005). Phosphorus is also involved in energy capture, storage, and transfer via ATP (Sterner and Elser, 2002). The input of “new” phosphorus to the oceans occurs primarily by geologically-slow orogeny, continental weathering, runoff, and possibly hydrothermal exchange (Vermeij, 1995; Tyrrell, 1999). Thus, “old” phosphorus sequestered in dead biomass (organic carbon) and sediments is recycled to sustain ecosystems as conditions (e.g., oxygen levels, sedimentation and bioturbation rates) permit (Dell'Anno and Danovaro, 2005). Despite variation in the C:P ratios of modern phytoplankton taxa (Table 1), fresh dead marine organic matter arriving at the sediment–water interface has approximately canonical (“average”) Redfield C:P ratios of ∼ 106:1 (Anderson et al., 2001). Early diagenesis initially causes C:P burial ratios of sediments to increase up to ∼ 400 in sediments younger than ∼ 5 Ma but most organic carbon is eventually remineralized and returned to the water column in dissolved form (i.e., buried TOC declines with increasing sediment age) while N 80% of phosphorus may remain in sediments in authigenic form or is precipitated as oxides. C:P ratios of sediments then decline through an additional ∼ 10 million years to a relatively constant value of ∼ 10:1 prior to ∼15 Ma (Anderson et al., 2001). Anderson et al. (2001) therefore concluded that the ratio of carbon-to-phosphorus buried in sediments should be used to assess the biogeochemical behavior of carbon and phosphorus in relation to primary productivity through geologic time rather than either TOC or phosphorus alone.

Fig. 2. Percent total organic carbon (TOC), %P, and C:P burial ratios. Y-axes: log10. X-axes: age in mega-annums (Ma). Note slight overlap of time scales in left and right columns. Solid black lines, least-squares fit including DSDP samples but excluding phosphorites and iron formations (see Methods). Black arrows in B, F indicate decreases and subsequent increases of C:P burial ratios from 2.0 to 1.0 Ga discussed in text. A, B) %TOC. Red diamonds: averages from Ronov (1976), not included in least-squares fits or statistical tests (Table 2); C, D) %P; E, F) C:P burial ratios. Orange line indicates 3-point moving average for all non-phosphorites for which both carbon and phosphorus data were available. C:P ratio of ∼ 3667:1 at 827 Ma also excluded from moving averages (see Methods).

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Our null hypothesis is therefore that C:P burial ratios older than ∼ 15 Ma would be expected to approximate 10:1 if nutrient levels, primary productivity, and carbon burial have remained constant. Alternatively, changes in the ratios may reflect significant shifts in food (phytoplankton) quantity (primary productivity) and quality (i.e., stoichiometry; Sterner and Elser, 2002) that are in turn related to the evolution of major phytoplankton taxa through geologic time. 3. Methods 3.1. C:P burial ratios C:P burial ratios were calculated by dividing %TOC by %P, as originally reported in the literature. If carbon and phosphorus were originally reported in other units (typically mg/g carbon/g sediment and P2O5, respectively), they were converted to %TOC and %P before use for the sake of standardization. C:P burial ratios were normally calculated only when both %TOC and %P were available for the same samples or stratigraphic unit from the same source, but in some cases calculations involved using data from different sources for the same stratigraphic units; these cases (∼25 data points) were almost exclusively for samples older than ∼2 Ga. Also, certain lithologies may be represented by only TOC or %P, but not both; thus, the number of samples for TOC exceeds the number of samples for %P, which in turn exceeds the number of samples for C:P burial ratios. Samples with likely terrestrial plant input (based on available sample descriptions or paleoenvironmental interpretations) were excluded. TOC, %P, and C:P burial ratios and sources are available from the senior author or from Elsevier Online. The use of percentages, ratios, or both can, conceivably, produce artifacts. Nevertheless, our methodology yields C:P burial ratios which approximate the Corganic/Preactive ratios obtained by the analytical techniques of Anderson et al. (2001) for Cenozoic marine sediments, where Preactive = phosphorus in all phases, i.e., Preactive = Porganic + Poxide-associated + Pauthigenic. Mort et al. (2007) used a similar approach in their study of Corg/Preact ratios during Cenomanian (early Late Cretaceous) Oceanic Anoxic Event (OAE) 2). Moreover, despite potential biases or “noise” introduced by different methodologies or units of measurement of various workers and their conversion to percentages, our methodology yields long-term trends that are corroborated by other indices and that are meaningful in the context of documented paleoenvironmental and paleobiological changes spanning geologic time. These indices are discussed below.

3.2. Age assignments of samples Age assignments for DSDP samples were corrected using published sediment accumulation curves for the respective sites or interpolation between stratigraphic boundaries (Leg 6, west central Pacific (Hawaii to Guam): sites 44, 47–50, 52–55, 59–60; Leg 39, north-tosouth Atlantic: sites 354–358; Leg 72, southwest Atlantic off South America: sites 515A–B, 516, 516F, 517–518; Leg 80, northeast Atlantic off Goban Spur: sites 549–550, 550B, 551; Leg 93, northwest Atlantic off U. S. midAtlantic coast: site 603B). Closely-spaced samples from other units for which accumulation rates were not available were all assigned to the same age, yielding a range of TOC, %P, and C:P burial ratios for the same age. Widely-spaced samples from units with coarse chronology were averaged and plotted as a single point for the unit. Otherwise, the chronology published by original authors was used unless revised ages were available from other sources (e.g., Gradstein et al., 2004). 3.3. Statistical analysis and curve fitting Data points were assigned to different lithologic groups based on the original description of the authors (Fig. 2). These plots indicated that although the TOC of phosphorites falls within the range of other lithologies, they also exhibit relatively low C:P burial ratios because of high %P (Fig. 2) that results from extreme diagenesis during phosphogenesis (Cook and McElhnny, 1979). Banded Iron Formations (BIFs) also result from extreme diagenesis that is thought to precipitate phosphorus onto Table 2 Comparison of Phanerozoic and Precambrian samples a

Phanerozoic Mean ± 95% Confidence Interval Median Range # Observations (n) Precambrian Mean ± 95% Confidence interval Median Range # Observations (n)

TOC (%)

P (%)

C:P burial ratios

2.04 ± 0.24

0.057 ± 0.007 53.2 ± 8.92

0.38 0.04 0.015–33.1 0.004–2.0 1036 634

8.9 0.1–956.8 634

0.8 ± 0.012

16.9 ± 5.66

0.08 ± 0.013

0.11 0.05 0.001–20.7 0.001–0.6 1578 216

4.75 0.02–141.7 122

Mann Whitney U (Phanerozoic versus Precambrian) U 5.2E5 5.9E4 2.7E4 p (same median) 2.7E-57 0.0018 2.6E-7 a

Phosphorites and Banded Iron Formations excluded (See Methods).

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iron oxides (Bjerrum and Canfield, 2002). Phosphorites and BIFs were therefore excluded from statistical analyses (Table 2), least-squares fits (Fig. 2), and other curve-fitting procedures (Fig. 3). The C:P burial ratio of

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3667:1 at 827 Ma was also excluded because it is an obvious outlier (Fig. 2). All other lithologies were used in analyses; these lithologies fall within the range of the chemical composition of normal shales, which

Fig. 3. Comparison of original (Fig. 2E and F) and smoothed C:P burial ratios with forg. Y-axes: log10. X-axes: age in mega-annums (Ma). Note changes in y-axis scales and slight overlap of time scales between left and right columns. Large black arrows in B, D, F indicate decreases and subsequent increases of C:P burial ratios from 2.0 to 1.0 Ga discussed in text. A, B) B-spline (decimation factor = 10). Similar results were obtained for decimation factors >10. A) Phanerozoic. B) Precambrian. C, D) Sinusoidal curves. C) Phanerozoic. First harmonic (dotted line): χ2 = 221.04, AIC = − 1.0442, R2 = 0.49782, p = 4.1725E-95; Second harmonic (solid line): χ2 = 283.81, AIC = −0.79428, R2 = 0.24867, p = 6.6879E-40; First and second harmonics combined (not shown): χ2 = 195.41, AIC = − 1.158, R2 = 0.48297, p = 1.2146E-88. D) Precambrian. First harmonic: χ2 = 95.344, AIC = − 0.19735, R2 = 0.083975, p = 0.010827. E, F) forg moving averages of Des Marais et al., 1992 (◆) and Hayes et al., 1999 (○).

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approximate the average chemical composition of sedimentary rocks over geologic spans of time, e.g., Paleozoic versus Meso-Cenozoic (Garrels and Mackenzie, 1971). Least squares fits are shown only to indicate the range of variation of the data. Percent TOC, %P, and C:P burial ratios each exhibited an approximately log-normal distribution for the Phanerozoic and Precambrian, respectively. Nonparametric procedures (Paleontological Statistics, or PAST, freeware; Hammer and Harper, 2006; PAST website: http://folk.uio.no/ohammer/past/fitting.html) were therefore used for statistical analysis (Table 2). Least-squares lines in Fig. 2A–F were fitted by standard procedures. Smoothed curves were fitted to the C:P burial ratios using several techniques for comparative purposes. We initially applied a 3-point moving average (Fig. 2E and F, orange curve), which uses the mean of the original value and its two neighbors; this resulted in very “spikey” curves because the data points exist at irregular intervals and/or multiple data points may exist for the same interval (Hammer, 2007, personal communication). We therefore fitted a B (cubic) spline (Fig. 3A and B), which is typically used to construct a smooth curve through noisy data sets by fitting a sequence of thirdorder polynomials (Hammer and Harper, 2006). In PAST, the number of data points contributing to each polynomial section is determined by a decimation factor set by the user; larger decimation yields a smoother curve. Data are then fitted with a least-squares criterion to the B-spline. The coefficients of each polynomial are selected so that segments of the B-spline are continuous, allowing the coefficients of each segment to be optimized so that they fit the data points within the respective interval as closely as possible (Hammer and Harper, 2006). The B-spline plots of this study (Fig. 3A and B) also display the original data points, at the recommendation of Hammer (2007, pers. comm.). Because of the apparent influence of tectonic cycles on phytoplankton stoichiometry, we tested for cyclicity using a sinusoidal curve-fitting function, which is useful for modeling periodicities in time series (Fig. 3C and D; Hammer and Harper, 2006). The algorithm used in the procedure is base on a least-squares criterion and singular value decomposition. Each sinusoid is given by: y ¼ a⁎ cos ð2⁎ p⁎ ð x  x0 Þ=T  pÞ; where a is the amplitude, T is the period, p is the phase, and x0 is the first (smallest) x value. The algorithm is very robust and almost guaranteed to find the global optimum (PAST website). In the default setting (used

here), the periods are initially set to the range of the x values, and harmonics (1/2, 1/3, etc. of the fundamental period). The Chi-squared (χ2) fitting error, coefficient of determination (R2, 0 = bad fit, 1 = perfect), p based on F test), and Akaike Information Criterion (AIC) are calculated, where: AIC ¼ 2k þ n ln ð RSS=nÞ; and k = number of parameters, n = number of observations, and RSS = residual sum of squares (Akaike, 1974). The AIC is a measure of the curve's goodness of fit (the equation assumes that the periods are estimated from the data; PAST website). Increasing the number of parameters improves the goodness of fit, but the AIC includes a penalty that is an increasing function of the number of parameters, so that overfitting is discouraged. Thus, the AIC should be as low as possible to maximize fit but avoid overfitting. 3.4. Bioturbation rates Bioturbation rates (Fig. 1) are from the compilations of Thayer (1983), who calculated rates from the chronostratigraphic ranges of higher level taxa known or believed to contain substantial numbers of bioturbators. Compilations included both extant and extinct taxa, thus countering the potential effects of differential preservation toward the present. Extrapolation of modern bioturbation rates to ancient representatives of the same taxa is supported by functional morphology. A curve for the number of taxa with mean feeding depth N10 cm is very similar to the bioturbation rate curve shown; although not a direct measure of reworking rate, the depth of bioturbation is preservable and may also be used to infer bioturbation rates through time (Thayer, 1983). Thayer's (1983; see also Vermeij, 1987) data were most recently used to examine predation and grazing through time by Kelley and Hansen (2001). 4. Results and discussion 4.1. The Phanerozoic The range of %TOC is quite broad (Fig. 2A and B) whereas, with the exception of phosphorites, %P tends to remain relatively constant and cluster around the least-squares fit (Fig. 2C and D). Both indices therefore appear to indicate relatively constant nutrient levels and primary productivity through time. Nevertheless, there is a statistically significant increase of TOC from the Proterozoic to the Phanerozoic, while %P appears to

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decrease somewhat between the two intervals (Table 2). Thus, C:P burial ratios of the Phanerozoic also tend to be higher (Table 2; Fig. 2E and F), falsifying the null hypothesis. Cyclic behavior of C:P burial ratios is also evident in the fitted curves. Despite its spikey nature (see Methods), the 3-point moving average broadly fluctuates around the 10:1 ratio during the Phanerozoic (Fig. 2E, compare 3-point moving average in orange with least-squares fit). The B-spline also indicates two broad cycles during the Paleozoic and Meso-Cenozoic, respectively (Fig. 3A; compare Fig. 2E). Despite the large fitting errors, the second harmonic of the sinusoidal fit behaves similarly, while the first harmonic indicates a broad rise into the early Mesozoic followed by a decline toward the present (Fig. 3C). The behavior of the B-spline and sinusoidal fits is corroborated by forg, the fraction of carbon buried as organic carbon (where forg = 1 − fcarb and fcarb = the fraction of carbon buried in carbonates; Fig. 3E; Hayes et al., 1999). C:P burial ratios would therefore appear to primarily reflect carbon burial, given the relative constancy of phosphorus levels through the Phanerozoic (Fig. 2C, least-squares fit) due to its geologically-slow input from other sources. High or rising C:P burial ratios and forg of the early-to-middle Paleozoic are, for example, associated with rifting continents and broad increases of sea level, warm climate (high CO2), sluggish ocean circulation, and anoxia, all of which would have presumably enhanced carbon burial irrespective of nutrient inputs (compare broad cycles of sea level in Fig. 1 with Fig. 3A and C). Nevertheless, other conditions associated with tectonic cycles also promoted carbon burial. Anoxia associated with supercontinent rifting during the early-tomiddle Paleozoic promoted phosphorus and iron release. Even so, the green phytoplankton of this time (“acritarchs”; Fig. 1) are thought to have diversified in open oligotrophic (nutrient-poor) waters overlying deeper anoxic ones (Martin, 1995, 1996, 2001, 2002; Quigg et al., 2003; Falkowski et al., 2004; Katz et al., 2004, 2005). These taxa were relatively small; the relatively small size of acritarchs decreased volume-to-surface (V/S) ratios which, in modern phytoplankton, increases the cell surface area available for nutrient absorption under limiting conditions (Hayes et al., 1999). It was during the early-and-middle Paleozoic that faunas with, presumably, low metabolic rates dominated. More active predators with higher metabolic rates, began to diversify in the Late Devonian (Fig. 1; Signor and Brett, 1984; Vermeij, 1987; Bambach, 1999). Also beginning in the Late Devonian was a broad rise of 87 Sr/86Sr isotope ratios that continued through the

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Permo-Carboniferous (Fig. 1). The broad rise of strontium isotopes suggest that nutrient inputs (e.g., phosphorus), and thus marine primary productivity, increased as a result of widespread orogeny, overall sealevel fall, and the spread of forests and greater rates of weathering (Fig. 1; cf. Tappan, 1968, 1971, 1986 and Pitrat, 1970; Martin, 1996, 2002, 2003). Enhanced deep-ocean circulation due to southern hemisphere glaciation also likely affected rates of nutrient upwelling and carbon burial. TOC, C:P burial ratios, and forg all eventually began to decrease later in the Permian, as phosphorite deposition, which depends on the upwelling of nutrient-rich waters resulting from the decay of organic matter (Cook and McElhinny, 1979), increased (Figs. 2 and 3). Although present during the Paleozoic, the Modern Fauna, which is characterized by active predation in diverse taxa, began to expand in earnest in the Mesozoic (Fig. 1). Organic matter burial also increased again during the Mesozoic and Paleogene in response to the initiation of a second cycle of supercontinent rifting and rising sea level (Figs. 1 and 2A; Fig. 3A and C). Unlike the early-tomiddle Paleozoic, however, nutrient runoff into widespread epeiric seas was probably enhanced by extensive terrestrial floras (Föllmi, 1995; Martin, 1996; Bambach, 1999; Katz et al., 2004, 2005; Lenton and Klausmeier, 2006). forg rose accordingly, implying higher rates of primary productivity (Fig. 3E; Hayes et al., 1999). Still, black shale deposition appears to have been much more episodic; peaks of black shale deposition in Figs. 2A and 3A correspond to oceanic anoxic events associated with phosphorus release and enhanced primary productivity (Mort et al., 2007). The episodic nature of these events suggests greater oxygenation of the oceans (as opposed to the early-to-middle Paleozoic) due presumably to carbon burial in the seas and on land (Falkowski et al., 2004; Katz et al., 2004, 2005). As oxygen levels rose, red phytoplankton lineages diversified (Quigg et al., 2003; Falkowski et al., 2004; Katz et al., 2004, 2005), culminating with the tremendous diversification of diatoms in the Neogene (Fig. 1). In summary, TOC would appear to primarily reflect long-term cycles of sea level that influenced carbon burial only, because %P remained relatively constant through the Phanerozoic (with the exception of phosphorites). However, the other conditions associated with these tectonic cycles (e.g., orogeny, oxygen levels) also influenced nutrient availability (including phosphorus), phytoplankton stocihiometry, and primary productivity through the Phanerozoic. The relative constancy of %P suggests chronic nutrient limitation that resulted from the geologically-slow input of nutrients and the sequestration

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of nutrients into ever larger amounts of biomass in the oceans and on land (Martin, 1996, 2002). Indeed, diatom frustule size decreased through the Neogene, suggesting nutrient limitation (Finkel et al., 2005), even as diatoms greatly diversified (Finkel et al., 2005). The behavior of TOC, C:P burial ratios, and forg from the post-Cambrian through the Eocene is consistent with the reworking (e.g., bioturbation) of phytoplankton products, given low rates of phytoplankton productivity, relatively high CO2 levels, and relatively small volume-to-surface (V/S) ratios (i.e. , nutrient limitation) of individual phytoplankton (Hayes et al., 1999). Increasing oxygen levels and the continued pelagic rain of dead organic matter (food) likely stimulated rates and depths of bioturbation and nutrient cycling, especially through the Meso-Cenozoic (Fig. 1; Thayer, 1983). TOC (especially of deep-sea sediments) began to decline by the Cenozoic (Fig. 2A, purple circles), well before sea-fall began in earnest in the midCenozoic, while %P of deep-sea sediments began to cluster more closely around the Phanerozoic trend (Fig. 2C, least-squares fit); thus, C:P ratios also began to decrease well before substantial sea-level fall began in the Cenozoic (Figs. 2E and 3A, C). In other words, carbon burial is not just a function of sea level. forg and δ13Ccarb finally followed suit about the midNeogene (∼ 15 Ma) (Figs. 1 and 3E), the timing of which corresponds to a further increase in the rates and depths of bioturbation (Fig. 1) and the diagenetic transformations described above for the most recent 15 Ma (Anderson et al., 2001). 4.2. The Precambrian The respective ranges of %TOC and C:P burial ratios for the Precambrian, like those of the Phanerozoic, are relatively broad (Fig. 2B and F). Also, with the exception of phosphorites (which are most prominent in the latest Proterozoic), %P tends to remain relatively constant (Fig. 2D, least-squares line). These indices therefore also appear to reflect relatively constant nutrient levels and primary productivity. However, a gradual, but nevertheless distinct, rise of C:P burial ratios began no later than ∼ 1.5–1.3 Ga (black arrows in Figs. 2F and 3B, D). The predominant lithology representing this trend is normal shales, which are considered to approximate the average chemical composition of sedimentary rocks over geologic scales of time (Garrels and Mackenzie, 1971). The behavior of C:P burial ratios from ∼ 2.0 to 1.0 Ga is corroborated by the behavior of TOC and forg, both of which declined after ∼ 2.0 Ga and then began to rise through the next

billion years (black arrows in Figs. 2B and 3F). The initial decline of forg after 2.0 Ga has been attributed to relatively low rates of tectonism and low-lying continents, which may have inhibited carbon burial (Des Marais et al., 1992; Brasier and Lindsay, 1998). Like the Phanerozoic, however, low rates of carbon burial may also reflect relatively low nutrient availability. Prior to ∼2 Ga, the oceans are thought to have been anoxic, making dissolved Fe available to ironbased nitrogenases used by cyanobacteria in nitrogen fixation (Anbar and Knoll, 2002; Katz et al., 2004, 2005). As oxygen levels slowly climbed into the Mesoproterozoic (due to cyanobacterial photosynthesis), phosphorus availability may have been limited by low-lying continents and erosion (Brasier and Lindsay, 1998) and adsorption of P onto iron oxide precipitates of Banded Iron Formations (Fig. 2D, red circles; Brasier and Lindsay, 1998; Anbar and Knoll, 2002; Bjerrum and Canfield, 2002). The initial rise of TOC, C:P burial ratios, and forg sometime after ∼2.0 Ga follows the appearance of probable acritarchs at 1.9–1.7 Ga (Butterfield, 1997, 2001), an increase of eukaryotic biomarkers (Jackson and Moore, 1976; Summons and Walter, 1990; Brocks et al., 2003), and overlaps with the divergence of green and red algal lineages sometime between 1.6 and 1.2 Ga, with the red line proliferating in coastal waters where oxygen was more prevalent (Anbar and Knoll, 2002; Katz et al., 2004, 2005). The gradual rise of these indices therefore suggests that nutrient limitation was beginning to ease somewhat, perhaps as a result of tectonism and increased weathering by rising oxygen levels (Brasier and Lindsay, 1998; Anbar and Knoll, 2002). These developments culminated in a distinct shift to more positive δ13C values during the Grenville Orogeny and the assembly of Rodinia by ∼ 1.25 Ga, which may have also influenced carbon burial (Fig. 1; Brasier and Lindsay, 1998; Anbar and Knoll, 2002; Katz et al., 2004, 2005). The subsequent wide swings of forg (Fig. 3F) during the Neoproterozoic appear associated with drastic climatic and paleoceanographic (“Snowball Earth”) regimes involving tectonism and glaciation (Hayes et al., 1999). There is some indication that C:P burial ratios may have fluctuated similarly (Fig. 2F, pink triangles indicating glaciogenic deposits at ∼ 827 Ma). After ~ 0.6 Ga, strontium isotope ratios (87Sr/86Sr) then increased to their highest values in Earth history during the Proterozoic-Phanerozoic transition, suggesting further nutrient runoff associated with widespread orogeny (Fig. 1; Montañez et al., 2000; Martin, 2002, 2003), while phosphorite deposition first became

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pronounced (Fig. 2D; Cook and McElhinny, 1979). The overall trend of increasing C:P burial ratios nevertheless continued toward the end of the Neoproterozoic, suggesting continued heightened primary productivity and carbon burial (Figs. 2F and 3B, D). This upward trend culminated in a significant upward shift of TOC and C:P burial ratios during the Proterozoic–Phanerozoic transition (Table 2). 5. Synthesis 5.1. Phytoplankton stoichiometry and the emergence of marine metazoans The fossil record indicates that the geologically-rapid diversification of marine metazoans began in the Late Neoproterozoic (ca. 600 Ma). The first definite traces of metazoans (shallow burrows) appeared no later than ∼ 560 Ma, followed by deeper burrowing and predation (Knoll et al., 2004). The evolution of larger metazoans prior to this time was presumably prohibited by their smaller surface-to-volume ratios, which limited oxygen diffusion (e.g., Falkowski et al., 2005; Fike et al., 2006; Canfield et al., 2006; Berner et al., 2007). Oxygen also appears to have influenced phytoplankton stoichiometry, as described above. Food quantity and, especially, food quality (autotroph stoichiometry) must in turn have been critical to early metazoan evolution; otherwise, the aerobic respiration of less nutritious substrates would have had less energetic (ATP) yield. Of the three basic consumer strategies that appeared during the Late Neoproterozoic (detritivores, herbivores, and carnivores), carnivores eat the most nutritionally-balanced foods (animals eating animals) and consistently exhibit the highest growth efficiencies and assimilation (i.e., conversion of food to consumer biomass), whereas detrivores (e.g., bioturbators) eat the least nutritious foods (animals eating non-living matter) and have low growth efficiencies and assimilation; herbivores occupy an intermediate position (Sterner and Elser, 2002). Bioturbation, locomotion, and, especially, predation are energetically expensive because they require the well-developed integration of nervous systems and musculatures (Thayer, 1983; Rhodes and Thayer, 1991; Rhodes and Thompson, 1993; Plotnick, 2006). We suggest that autotroph stoichiometry evolved in parallel with the oxygenation of the atmosphere and oceans and, like oxygenation, began much earlier during the Proterozoic, long before the appearance of bona fide metazoans or their traces in the fossil record. Beginning ∼ 1.5 Ga, Proterozoic marine ecosystems resembled

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modern “stoichiometrically-imbalanced” ecosystems, in which eukaryotic autotrophs fix large amounts of carbon despite low dissolved nutrient levels (Sterner and Elser, 2002). Under imbalanced conditions in modern ecosystems, nutrient use efficiency (= rate of C fixed/unit nutrient) of autotrophs is high but carbon use efficiency (= secondary production/primary production) by herbivores is low because the food is nutrient-poor (carbonrich) and excess carbon must be eliminated by animals to obtain nutrients necessary for growth (Sterner and Elser, 2002). This results in the reduced importance of herbivores as carbon conduits in food webs, leaving more carbon available for detritivores (e.g., bioturbators), which have low net growth efficiencies; decomposition rates are therefore low and carbon burial is favored (Sterner and Elser, 2002). These conditions account for the gradual rise of TOC, C:P burial ratios, and forg beginning at ∼ 1.5 Ga (black arrows in Figs. 2B, F and 3B, D, F). The appearance of herbivores and carnivores in the fossil record during the significant upward shift of TOC and C:P burial ratios (Table 2) indicates that marine ecosystems were transiting to a “stoichiometricallybalanced” state during the Proterozoic–Phanerozoic transition (Sterner and Elser, 2002). Under these conditions in modern ecosystems, phytoplankton produce biomass of approximately Redfield proportions under more-or-less nutrient-saturated conditions; this biomass is capable of consistently sustaining higher trophic levels (Sterner and Elser, 2002). A possible modern analog for these ancient events is the eastern Mediterranean (Thingstad et al., 2005). Here, the experimental input of phosphorus to ultraoligotrophic, phosphorus-starved surface waters stimulates rapid luxury consumption of phosphorus by eukaryotic phytoplankton and bacteria. Thus, concentrations of intracellular nutrients of these cells increase above those immediately required for growth, making them more nutrient-rich (Sterner and Elser, 2002); phosphorus is therefore “tunneled” into mesozooplankton populations much more rapidly than through a typical food chain (Thingstad et al., 2005). A similar mechanism was postulated to have occurred during the NeoproterozoicPhanerozoic transition by Martin (1996, 2002; Fig. 1). TOC, C:P burial ratios, and forg tended to level off or decline somewhat after the significant upward shift of indices (Figs. 2E–F and 3), as %P declined somewhat (Table 2). We suggest that massive sequestration of nutrients into biomass, and thus severe nutrient limitation, occurred during this phase (Zhuravlev, 2001). Organisms of higher trophic levels tend to become enriched in phosphorus because of the widespread

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phosphorus-dependent energy transfer and information storage pathways involved in protein synthesis (Sterner and Elser, 2002). Other lines of evidence support the inference of severe nutrient limitation. First, the Late Neoproterozoic–Phanerozoic transition is when phosphorite deposition, which depends on the remineralization of organic matter (Cook and McElhinny, 1979), with its sequestered nutrients, first became pronounced (Fig. 2C and D). Second, acritarchs averaged ∼ 20 μm in diameter, which falls toward the lower limits of eukaryotic microplankton (range: 20–200 μm) and nannoplankton (2–20 μm) of the Meso-Cenozoic (Huntley et al., 2006); the extremely small size of acritarchs suggests adaptation to nutrient-poor conditions (Martin, 1996; see also Hayes et al., 1999). Third, C:P burial ratios and forg both declined into the early Phanerozoic, as depths of bioturbation increased (Thayer, 1983; Droser and Bottjer, 1988, 1989; Droser et al., 1999; Bottjer et al., 2000). Once triggered, rapid biosphere expansion would have necessitated the onset of nutrient recycling to sustain biosphere metabolism on ecologic time scales and biosphere complexification on geologic ones. The succession of fossil evidence during the Proterozoic–Phanerozoic transition has previously been interpreted to indicate a trophic cascade-like process like that of lacustrine ecosystems, but prolonged over geologic scales of time (Butterfield, 1997, 2001). Modern trophic cascades are said to result primarily from “top-down” control of community structure by upper trophic levels (herbivorous mesozooplankton and predators) that affect multiple lower trophic levels (Sterner and Elser, 2002). However, herbivorous mesozooplankton involved in modern cascades are fast-growing generalists with high nutrient demands and thus are strongly constrained by “bottom-up” autotroph stoichiometry, i.e., Redfield C:P ratios of autotroph biomass must be sufficient to sustain zooplankton metabolism (cf. Butterfield, 1997, 2001, and Sterner and Elser, 2002). Cropping by mesozooplankton would have increased nutrient cycling, thereby maintaining autotroph populations in a state of more rapid growth and increasing their nutrient content, i.e., food quality would have increased in the water column, allowing additional trophic levels to be added while lowering biomass yield (Sterner and Elser, 2002) and simultaneously increasing the rain of organic matter leaking to the benthos (Tyson, 1995; Logan et al., 1995). Organic matter decomposition and nutrient cycling were therefore favored but carbon burial was not (Sterner and Elser, 2002), as evidenced by the decline of C:P burial ratios and forg into the Cambrian.

5.2. Implications for biodiversification during the Phanerozoic The approximate C:P burial ratio of 10:1 reached during the Proterozoic–Phanerozoic transition suggests that fresh organic matter with approximately canonical Redfield ratios was being more-or-less consistently produced by phytoplankton, i.e., stoichiometricallybalanced conditions capable of supporting higher trophic levels were reached during this time. The Paleozoic is nevertheless characterized by relatively low and constant biodiversity and low rates of bioturbation (Fig. 1); moreover, the phytoplankton of the Paleozoic appear to have been adapted to nutrient-poor conditions and were themselves nutrient-poor (i.e., carbon-rich) as compared to later red lineages (Table 2). Thus, the growth efficiency of early-to-middle Paleozoic metazoans may have remained limited because they had to expend energy to burn off “excess” carbon in food before the associated nutrients could be used for biomass growth. Like the Proterozoic–Phanerozoic transition (Section 5.1), this nutrient imbalance may have delayed the appearance of more advanced carnivores until the PermoCarboniferous. Widespread orogeny, a general fall of sea level, and the spread of forests and greater weathering all occurred during this interval (Fig. 1). Enhanced deepocean circulation due to southern hemisphere glaciation also likely increased rates of upwelling and oxygenation. Increasing oxygenation of the oceans during the Permo-Carboniferous may have set the stage for the diversification of red phytoplankton lineages following the end-Permian extinctions. Black shales, for example, are much less common during the Mesozoic than the early-to-middle Paleozoic (Figs. 2 and 3). Also, the diversity and abundance of acritarchs, which belong to the green phytoplankton lineage, greatly declined and never recovered after about the Late Devonian (Tappan, 1968, 1971, 1980; Katz et al., 2004, 2005); the decline may therefore be related to greater nutrient inputs and the evolving stoichiometry of the oceans (Martin, 1996). The subsequent diversification of the Modern Fauna during the Meso-Cenozoic may consequently reflect the appearance of food quantity and quality sufficient to enhance metabolism and biomass during this time. Indeed, many of the crown metazoan taxa of the Modern Fauna that appeared during the “Cambrian explosion” did not begin to diversify substantially until the Mesozoic. The concomitant increases in the rates and depths of bioturbation through the Phanerozoic (Fig. 1) therefore suggest that, rather than simply reflecting a “passive” invasion of new habitats by an expanding biosphere, increasing bioturbation and nutrient cycling

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were necessary to sustain primary productivity and energetics because of the geologically-slow input of macronutrients like phosphorus from land and the continued sequestration of nutrients into marine and terrestrial biomass. 6. Conclusions Diversification of the marine biosphere is intimately linked to the evolution of the biogeochemical cycles of carbon, nutrients, and primary productivity. Food quantity and, particularly, food quality (autotroph stoichiometry) were critical to early metazoan evolution because the aerobic respiration of less nutritious substrates would have yielded less energy and been unable to sustain trophic levels higher than detritivores. Evolving food quantity and quality was primarily a function of broad tectonic cycles that influenced carbon burial not just through sea-level change but also via nutrient inputs from land and sea. Beginning no later than ∼1.5 Ga, the marine biosphere began to transition from “stoichiometrically-imbalanced” ecosystems, in which eukaryotic autotrophs fix large amounts of carbon despite low dissolved nutrient levels, to “stoichiometrically-balanced” ones, in which phytoplankton produce biomass of approximately Redfield proportions under more-or-less nutrient-saturated conditions. This food source was finally capable of sustaining higher trophic levels. The resulting geologically-rapid diversification of the marine biosphere sequestered nutrients into increasing amounts of biomass, triggering severe nutrient limitation and the onset of increasing rates and depths of bioturbation and nutrient cycling. Carbonrich (i.e., relatively nutrient-poor) green phytoplankton lineages nevertheless continued to dominate through at least the early-to-middle Paleozoic, paralleling the dominance of (apparently) less energetic and less diverse Cambrian and Paleozoic macrofaunas. The growth efficiency of early-to-middle Paleozoic metazoans was therefore limited because they had to expend energy to respire excess carbon in their food before the associated nutrients could be used for growth. The subsequent diversification of the Modern Fauna during the Meso-Cenozoic was paralleled by the diversification of more nutrient-rich (carbon-poor) red phytoplankton lineages. The diversification of MesoCenozoic marine faunas therefore appears to reflect the appearance of food quantity and quality sufficient to further accelerate metabolism and growth (biomass increase). The concomitant increases in the rates and depths of bioturbation through the Phanerozoic thus suggest that, rather than simply reflecting a “passive”

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invasion of new habitats by an expanding biosphere, increasing bioturbation and nutrient cycling were necessary to sustain primary productivity and energetics because of the geologically-slow input of macronutrients like phosphorus from land and the continued sequestration of nutrients into marine and terrestrial biomass. Increasing primary productivity and, especially, nutrient cycling therefore appear to have been critical to sustaining the expansion and diversification of the marine biosphere over time spans of many tens to hundreds-of-millions of years. Acknowledgements We thank an anonymous reviewer but, especially, Wolfgang Kiessling (Berlin) for his critical, constructive, and supportive criticism of the original and revised manuscripts. We also owe sincere thanks to Øyvind Hammer (Oslo) for his statistical advice and the modification of the PAST cubic spline freeware to make it suitable for use with our data, and to Togwell Jackson (Winnipeg) for his initial advice and encouragement. Our thanks also to Mimi Katz for artwork used in portions of Fig. 1. This paper is an outgrowth of taphonomic research supported by the National Science Foundation. A much earlier version of the manuscript was presented by REM at the International Conference on Taphonomy, Valencia, Spain (Martin, 2002). REM thanks the organizers, especially Miquel De Renzi, for the invitation to present a keynote address at the conference and for their warmth, hospitality, and collegiality. The University of Delaware Department of Geological Sciences kindly defrayed the costs of the color reproduction of Figures 1 and 2. References Aberhan, M., Kiessling, W., Fürsich, F.T., 2006. Testing the role of biological interactions in the evolution of mid-Mesozoic marine benthic ecosystems. Paleobiology 32, 259–277. Akaike, H., 1974. A new look at the statistical model identification. IEEE Trans. Automat. Contr. 19, 716–723. Allmon, W.D., Ross, R.M., 2001. Nutrients and evolution in the marine realm. In: Allmon, W.D., Bottjer, D.J. (Eds.), Evolutionary Paleoecology: The Ecological Context of Macroevolutionary Change. Columbia Univ. Press, NY, pp. 105–148. Alroy, J., Marshall, C.R., Bambach, R.K., Bezusko, K., Foote, M., Fürsich, F.T., Hansen, T.A., Holland, S.M., Ivany, L.C., Jablonski, D., Jacobs, D.K., Jones, D.C., Kosnik, M.A., Lidgard, S., Low, S., Miller, A.I., Novack-Gottshall, P.M., Olszewski, T.D., Patzkowsky, M.E., Raup, D.M., Roy, K., Sepkoski, J.J., Sommers, M.G., Wagner, P.J., Webber, A., 2001. Effects of sampling standardization on estimates of Phanerozoic marine diversification. Proc. Natl. Acad. Sci. 98, 6261–6266.

290

R.E. Martin et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 277–291

Anbar, A.D., Knoll, A.H., 2002. Proterozoic ocean chemistry and evolution: a bioinorganic bridge? Science 297, 1137–1142. Anderson, L.D., Delaney, M.L., Faul, K.L., 2001. Carbon to phosphorus ratios in sediments: implications for nutrient cycling. Glob. Biogeochem. Cycles 15, 65–79. Bambach, R.K., 1993. Seafood through time: changes in biomass, energetics, and productivity in the marine ecosystem. Paleobiology 19, 372–397. Bambach, R.K., 1999. Energetics in the global marine fauna: a connection between terrestrial diversification and change in the marine biosphere. GeoBios 32, 131–144. Bambach, R.K., Knoll, A.H., Sepkoski, J.J., 2002. Anatomical and ecological constraints on Phanerozoic animal diversity in the marine realm. Proce. Natl. Acad. Sci. 99, 6854–6859. Berner, R.A., VandenBrooks, J.M., Ward, P.D., 2007. Oxygen and evolution. Science 316, 557–558. Bjerrum, C.J., Canfield, D.E., 2002. Ocean productivity before about 1.9 Gyr ago limited by phosphorus adsorption onto iron oxides. Nature 417, 159–162. Bottjer, D.J., Hagadorn, J.W., Dornbos, S.Q., 2000. The Cambrian substrate revolution. GSA Today 10, 1–7. Brasier, M.D., Lindsay, J.F., 1998. A billion years of environmental stability and the emergence of eukaryotes: new data from northern Australia. Geology 26, 555–558. Brocks, J.J., Buick, R., Summons, R.E., Logan, G.A., 2003. A reconstruction of Archean biological diversity based on molecular fossils from the 2.78 to 2.45 billion-year-old Mount Bruce Supergroup, Hamersley Basin, Western Australia. Geochim. Cosmochim. Acta 67, 4321–4335. Butterfield, N.J., 1997. Plankton ecology and the Proterozoic– Phanerozoic transition. Paleobiology 23, 247–262. Butterfield, N.J., 2001. Ecology and evolution of Cambrian plankton. In: Zhuravlev, A. Yu., Riding, R. (Eds.), The Ecology of the Cambrian Radiation. Columbia Univ. Press, New York, NY, pp. 200–216. Butterfield, N.J., 2004. A vaucheriacean alga from the Middle Neoproterozoic of Spitsbergen: implications for the evolution of Proterozoic eukaryotes and the Cambrian explosion. Paleobiology 30, 231–252. Canfield, D.E., Poulton, S.W., Narbonne, G.M., 2006. Late-Neoproterozoic deep-ocean oxygenation and the rise of animal life. Science 315, 92–95. Cook, P.J., McElhinny, M.W., 1979. A reevaluation of the spatial and temporal distribution of sedimentary phosphate deposits in the light of plate tectonics. Econ. Geol. 74, 315–330. Dell'Anno, A., Danovaro, R., 2005. Extracellular DNA plays a key role in deep-sea ecosystem functioning. Science 309, 2179. Des Marais, D.J., Strauss, H., Summons, R.E., 1992. Carbon isotope evidence for the stepwise oxidation of the Proterozoic environment. Nature 359, 605–609. Droser, M.L., Bottjer, D.J., 1988. Trends in depth and extent of bioturbation in Cambrian carbonate marine environments, Western United States. Geology 16, 233–236. Droser, M.L., Bottjer, D.J., 1989. Ordovician increase in extent and depth of bioturbation; implications for understanding early Paleozoic ecospace utilization. Geology 17, 850–852. Droser, M.L., Gehling, J.G., Jensen, S., 1999. When the worm turned: concordance of Early Cambrian ichnofabric and trace-fossil record in siliciclastic rocks of South Australia. Geology 27, 625–628. Falkowski, P.G., Katz, M.E., Knoll, A.H., Quigg, A., Raven, J.A., 2004. The evolution of modern eukaryotic phytoplankton. Science 305, 354–360.

Falkowski, P.G., Katz, M.E., Milligan, A.J., Fennel, K., Cramer, B.S., Aubry, M.P., Berner, R.A., Novacek, M.J., Zapol, W.M., 2005. The rise of oxygen over the past 205 million years and the evolution of large placental mammals. Science 309, 2202–2204. Fensome, R.A., Saldarriaga, J.F., Taylor, F.J.R. “Max”, 1999. Dinoflagellate phylogeny revisited: reconciling morphological and molecular based phylogenies. Grana 38, 66–80. Fike, D.A., Grotzinger, J.P., Pratt, L.M., Summons, R.E., 2006. Oxidation of the Ediacaran ocean. Nature 444, 744–747. Finkel, Z.V., Katz, M.E., Wright, J.D., Schofield, O.M.E., Falkowski, P.G., 2005. Climatically driven macroevolutionary patterns in the size of marine diatoms over the Cenozoic. Proc. Natl. Acad. Sci. U. S. A. 102, 8927–8932. Föllmi, K.B., 1995. 160 m.y. record of marine sedimentary phosphorus burial: coupling of climate and continental weathering under greenhouse and icehouse conditions. Geology 23, 859–862. Garrels, R.M., Mackenzie, F.T., 1971. Evolution of Sedimentary Rocks. Norton, New York. 397 pp. Gradstein, F.M., Ogg, J.G., Smith, A.G. (Eds.), 2004. A Geologic Time Scale. Cambridge University Press, Cambridge, UK, p. 589. Hammer, Ø., Harper, D., 2006. Paleontological Data Analysis. Blackwell, Oxford, UK. 351 pp. Hayes, J.M., Strauss, H., Kaufman, A.J., 1999. The abundance of 13C in marine organic matter and isotopic fractionation in the global biogeochemical cycle of carbon during the past 800 Ma. Chem. Geol. 161, 103–125. Huntley, J.W., Xiao, S., Kowalewski, M., 2006. 1.3 Billion years of acritarch history: an empirical morphospace approach. Precambrian Res. 144, 52–68. Jablonski, D., Roy, K., Valentine, J.W., Price, R.M., Anderson, P.S., 2003. The impact of the pull of the recent on the history of marine diversity. Science 300, 1133–1135. Jackson, T.A., Moore, C.B., 1976. Secular variations in kerogen structure and carbon, nitrogen and phosphorus concentrations in prePhanerozoic and Phanerozoic sedimentary rocks. Chem. Geol. 18, 107–136. Katz, M.E., Finkel, Z.V., Grzebyk, D., Knoll, A.H., Falkowski, P.G., 2004. Evolutionary trajectories and biogeochemical impacts of marine eukaryotic phytoplankton. Annu. Rev. Ecol. Evol. Syst. 35, 523–556. Katz, M.E., Wright, J.D., Miller, K.G., Cramer, B.S., Fennel, K., 2005. Biological overprint of the geological carbon cycle. Mar. Geol. 217, 323–338. Kelley, P.H., Hansen, T., 2001. Mesozoic marine revolution. In: Briggs, D.E.G., Crowther, P.R. (Eds.), Paleobiology II: A Synthesis. Blackwell, Oxford, UK, pp. 94–97. Kidwell, S.M., 2005. Shell composition has no net impact on largescale evolutionary patterns in molluscs. Science 307, 914–917. Kidwell, S.M., Brenchley, P.J., 1994. Patterns of bioclastic accumulation through the Phanerozoic: changes in input or destruction? Geology 22, 1139–1143. Kilham, P., Kilham, S.S., 1980. The evolutionary ecology of phytoplankton. In: Morris, I. (Ed.), The Physiological Ecology of Phytoplankton. Univ. California Press, Berkeley, pp. 571–597. Knoll, A.H., Walter, M.R., Narbonne, G.M., Christie-Blick, N., 2004. A new period for the geologic time scale. Science 305, 621–622. Kowalewski, M., Kiessling, W., Aberhan, M., Fürsich, F.T., Scarponi, D., Wood, S.L.B., Hoffmeister, A.P., 2006. Ecological, taxonomic, and taphonomic components of the post-Paleozoic increase in sample-level species diversity of marine benthos. Paleobiology 32, 533–561.

R.E. Martin et al. / Palaeogeography, Palaeoclimatology, Palaeoecology 258 (2008) 277–291 Lenton, T.M., Klausmeier, C.A., 2006. Co-evolution of phytoplankton C:N:P stoichiometry and the deep ocean N:P ratio. Biogeosciences Discuss. 3, 1023–1047. Logan, G.A., Hayes, J.M., Hieshima, G.B., Summons, R.E., 1995. Terminal Proterozoic reorganization of biogeochemical cycles. Nature 376, 53–56. Madin, J.S., Alroy, J., Aberhan, M., Fürsich, F.T., Kiessling, W., Kosnik, M.A., Wagner, P.J., 2006. Statistical independence of escalatory ecological trends in Phanerozoic marine invertebrates. Science 312, 897–900. Martin, R.E., 1995. Cyclic and secular variation in microfossil biomineralization: clues to the biogeochemical evolution of Phanerozoic oceans. Glob. Planet. Change 11, 1–23. Martin, R.E., 1996. Secular increase in nutrient levels through the Phanerozoic: implications for productivity, biomass, and diversity of the marine biosphere. Palaios 11, 209–219. Martin, R.E., 1998a. One Long Experiment: Scale and Process in Earth History. Columbia Univ. Press, New York, NY. 262 pp. Martin, R.E., 1998b. Catastrophic fluctuations in nutrient levels as an agent of mass extinction: upward scaling of ecological processes? In: McKinney, M.L., Drake, J.A. (Eds.), Biodiversity Dynamics: Turnover of Populations, Taxa, and Communities. Columbia Univ. Press, NY, pp. 405–429. Martin, R.E., 1999. Taphonomy: A Process Approach. Cambridge Univ. Press, Cambridge, UK. 508 pp. Martin, R.E., 2001. Marine plankton. In: Briggs, D.E.G., Crowther, P.R. (Eds.), Paleobiology II: A Synthesis. Blackwell, Oxford, UK, pp. 309–312. Martin, R.E., 2002. Cyclic and secular trends in preservation through geologic time: implications for the evolution of biogeochemical cycles. In: De Renzi, M., Alonso, M.V.P., Belinchón, M., Peñalver, E., Montoya, P., Márquez-Aliaga, A. (Eds.), Current Topics on Taphonomy and Fossilization—Taphos: International Conference on Taphonomy. Ayuntamiento de Valencia. Valencia, Spain, pp. 67–76. Martin, R.E., 2003. The fossil record of biodiversity: nutrients, productivity, habitat area and differential preservation. Lethaia 36, 179–193. Martin, R.E., Quigg, A., Podkovyrov, V., 2006. The evolution of ocean stoichiometry and diversification of the marine biosphere. Geol. Soc. America Abstracts with Programs 38 (7), 148. Montañez, I.P., Osleger, D.A., Banner, J.L., Mack, L.E., 2000. Evolution of the Sr and C isotope composition of Cambrian oceans. GSA Today 10, 1–7. Mort, H.P., Adatte, T., Föllmi, K.B., Keller, G., Steinmann, P., Matera, V., Berner, Z., Stüben, D., 2007. Phosphorus and the roles of productivity and nutrient recycling during oceanic anoxic event 2. Geology 35, 483–486. Peters, S.E., Foote, M., 2001. Biodiversity in the Phanerozoic: a reinterpretation. Paleobiology 27, 583–601. Pitrat, C.W., 1970. Phytoplankton and the late Paleozoic wave of extinction. Palaeogeogr., Palaeoclimatol., Palaeoecol. 8, 49–55. Plotnick, R.E., 2006. Chemoreception, odor landscapes, and foraging during the Ediacaran–Cambrian. Geol. Soc. America Abstracts with Programs 38 (7), 475. Quigg, A., Finkel, Z.V., Irwin, A.J., Rosenthal, Y., Ho, T.-Y., 2003. The evolutionary inheritance of elemental stoichiometry in marine phytoplankton. Nature 425, 291–294. Raup, D.M., 1976. Species diversity in the Phanerozoic: an interpretation. Paleobiology 2, 289–297. Raup, D.M., 1979. Biases in the fossil record of species and genera. Bull. Carnegie Mus. Nat. Hist. 13, 85–91. Rhodes, M.C., Thayer, C.W., 1991. Mass extinctions: ecological selectivity and primary production. Geology 19, 877–880.

291

Rhodes, M.C., Thompson, R.J., 1993. Comparative physiology of suspension-feeding in living brachiopods and bivalves: evolutionary implications. Paleobiology 19, 322–334. Ronov, A.B., 1976. Global carbon geochemistry, volcanism, carbonate accumulation, and life. Geochem. Int. 13, 172–195. Schidlowski, M., Hayes, J.M., Kaplan, I.R., 1983. Isotopic inferences of ancient biochemistries: carbon, sulfur, nitrogen, and hydrogen. In: Schopf, J.W. (Ed.), Earth's Earliest Biosphere: Its Origin and Evolution. Princeton Univ. Press, Princeton, NJ, pp. 149–186. Sepkoski, J.J., Bambach, R.K., Droser, M.L., 1991. Secular changes in Phanerozoic event bedding and the biological overprint. In: Einsele, G., Ricken, W., Seilacher, A. (Eds.), Cycles and Events in Stratigraphy. Springer Verlag, Berlin, pp. 298–312. Signor, P.W., Brett, C.E., 1984. The mid-Paleozoic precursor to the Mesozoic marine revolution. Paleobiology 10, 229–245. Sterner, R.W., Elser, J.J., 2002. Ecological Stoichiometry: The Biology of Elements from Molecules to the Biosphere. Princeton Univ. Press, Princeton, NJ. 439 pp. Strauss, H., Des Marais, D.J., Hayes, J.M., Summons, R.E., 1992. Concentrations of organic carbon and maturities and elemental compositions of kerogens. In: Schopf, J.W., Klein, C. (Eds.), The Proterozoic Biosphere. Cambridge Univ. Press, Cambridge, UK, pp. 95–99. Summons, R.E., Walter, M.R., 1990. Molecular fossils and microfossils of prokaryotes and protists from Proterozoic sediments. Am. J. Sci. 290–A, 212–244. Tappan, H., 1968. Primary production, isotopes, extinctions and the atmosphere. Palaeogeog., Palaeoclimatol., Palaeoecol. 4, 187–210. Tappan, H., 1971. Microplankton, ecological succession and evolution. Proceedings North American Paleontological Convention, part H, pp. 1058–1103. Tappan, H., 1980. Paleobiology of Plant Protists. W.H. Freeman, San Francisco, CA. Tappan, H., 1986. Phytoplankton: below the salt at the global table. J. Paleontol. 60, 545–554. Thayer, C.W., 1983. Sediment-mediated biological disturbance and the evolution of marine benthos. In: Tevesz, M.J.S., McCall, P.L. (Eds.), Biotic Interactions in Recent and Fossil Benthic Communities. Plenum Press, New York, NY, pp. 479–625. Thingstad, T.F., Krom, M.D., Mantoura, R.F.C., Flaten, G.A.F., Groom, S., 2005. Nature of phosphorus limitation in the ultraoligotrophic eastern Mediterranean. Science 309, 1068–1071. Tyrrell, T., 1999. The relative influences of nitrogen and phosphorus on oceanic primary production. Nature 400, 525–531. Tyson, R.V., 1995. Sedimentary Organic Matter: Organic Facies and Palynofacies. Chapman and Hall, London. 615 pp. Vermeij, G.J., 1987. Evolution and Escalation. Princeton University Press, Princeton, NJ. Vermeij, G.J., 1995. Economics, volcanoes, and Phanerozoic revolutions. Paleobiology 21, 125–152. Wagner, P.J., Kosnik, M.A., Lidgard, S., 2006. Abundance distributions imply elevated complexity of post-Paleozoic marine ecosystems. Science 314, 1289–1292. Zhuravlev, A. Yu., 2001. Biotic diversity and structure during the Neoproterozoic–Ordovician transition. In: Zhuravlev, A. Yu., Riding, R. (Eds.), The Ecology of the Cambrian Radiation. Columbia Univ. Press, New York, NY, pp. 173–199. Zotin, A.I., 1984. Bioenergetic trends of evolutionary progress of organisms. In: Lamprecht, I., Zotin, A.I. (Eds.), Thermodynamics and Regulation of Biological Processes. Walter de Gruyter, Berlin, pp. 451–458.