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FIGURE 8-7.

Scenario for early biotic evolution. (After Gogarten-Boekels et al. 1995.)

limit for growth, a more plausible limit on emergence than the optimal temperature, is used. There are now over a dozen organisms whose genomes have been completely sequenced, with more to come (Doolittle 1998b). The results are now confounding molecular phylogenists (Pennisi 1998). Horizontal gene transfer between coexisting organisms appears now to have been the rule rather than the exception (Olendzanski and Gogarten 1998), which makes the interpretation of phylogenies from even complete genomes difficult. Nevertheless, the original hypothesis ofWoese (1987) that early life was thermophilic or even hyperthermophilic (Pace 1997; Barns and Nierzwicki-Bauer 1997) is still robust (Gogarten, personal communication). Wachtershauser (1998) argues persuasively for a high temperature origin of life and by implication for a thermophilic LCA from the high improbability of enzymatic systems adapting from mesophily to hyperthermophily rather than the reverse path.

Surface cooling of the biosphere has been determined by its evolution, constrained by abiotic boundary conditions (i.e., luminosity of the Sun, continental area, outgassing rate). We proposed that soon after the origin of life this process commenced with the colonization of land by thermophiles/hy-

perthermophiles at around 3.8 Ga, resulting in the enhancement of weathering rates, and the sequestering of carbon dioxide into limestone deposits, thereby cooling the Earth's surface (Schwartzman and Volk 1989, 1991a; Oberbeck and Mancinelli 1994). Natural selection should have promoted the survival ofthose mutants among land biota with greater nutrient (essential minor and trace elements) extraction ability and water retention growth habits. The binding of mineral and rock particles by extracellular polymers such as polysaccharides produced by procaryotic colonies is well known from studies of natural consortia growing in and on the surface of soils (Campbell 1979). These characteristics are precisely those that would have increased the weathering rate per land area, thus accelerating the removal of CO2, bring down global temperatures.

Could higher levels of ultraviolet (UV) radiation at the surface have prevented microbial colonization of land in the Archean? Other gases in the atmosphere such as traces of sulfur dioxide could have acted as a UV screen in the Archean (Kasting et al. 1989), in the absence of oxygen and hence an ozone screen. Even without other UV screens, much higher levels of UV radiation at the surface could have been shielded by the organic products (e.g., mucus) of the bacterial colonies themselves (Lovelock 1988) and the stro-matolitic growth habit (Margulis et al. 1976) (but for another view on this issue see Towe 1994, who contends that minor levels of free oxygen provided just this UV screen). Pierson (1994) argued that an ozone screen was likely a necessary condition for large, multicellular life on land. Thus, the rise ofatmospheric oxygen after about 2 Ga, and its concomitant strong UV shield may be linked to a significant increase in the productivity of the land biota, particularly for eucaryotic algae, and biotic enhancement of weathering, leading to the apparent profound cooling in the mid-Proterozoic.

Why did atmospheric oxygen rise in the Proterozoic if oxygenic photosynthesis began at least 1.3 billion years earlier? What was the "sink" for the oxygen produced by cyanobacteria? The oxidation of dissolved ferrous iron in the ocean was likely a major sink. Another likely contributing sink was the reaction of reduced gases, such as hydrogen, from volcanic outgassing. Kasting, Eggler, and Raeburn (1993) argued that the key factor in the rise of oxygen was the progressive oxidation of the Archean mantle from the subduction of water in the hydrated sea floor and the release of reduced volcanic gases. By about 2 Ga, the upper mantle was oxidized enough so the oxidation state of volcanic gases increased, leading to a decrease in the consumption rate of atmospheric oxygen.

TABLE 8-2. Evolution of Land Biota

So il Commun ity (New or Dom inant Land Biota) Hyperthermophiles, thermophiles (methanogens, thermoacidophiles,

green nonsulphur bacteria, Thermus) Cyanobacteria

Cyanobacteria mats with eucaryotesa Eucaryotic/procaryotic mats'

Algal mats, primitive lichens (primitive fungi?), first Metazoa Bryophytesc (the emergence of Hypersea; see McMenamin and

McMenamin (1994) Vascular plants, lichensd Vascular plants (angiosperms, grasses)

aThis likely community can be viewed as an anti-lichen by the reversal of the procaryote/ eucaryote spatial relationship (i.e., procaryotic matrix with embedded eucaryotic cells, the inverse of lichen with cyanobacteria as the phycobiont). Perhaps modern microbial mats (see description of species diversity in Brown et al. 1985) are in some ways a model of ancient anti-lichens.

bFor example, possibly the actinolichen, a symbiosis of actinobacteria and green algae. Actinolichens have been synthesized in the laboratory, and one example has been reported in nature (Hawksworth 1988).

cEarliest fossil: lower middle Ordovician (Strother et al. 1996).

dEarliest fossil of vascular plant, early Silurian (Cai et al. 1996); earliest fossil of lichen, early Devonian (Taylor et al. 1995).

Each new innovation in the microbial soil community resulted in greater biotic enhancement of weathering, culminating in the rhizosphere of higher plants (Schwartzman et al. 1993). The long-term increase of a major factor— soil stabilization—as biotic evolution changed the dominant land biota from procaryotes to vascular plants in the Phanerozoic has already been discussed. A somewhat speculative history is shown in table 8-2 (at this stage it must be speculative because the fossil record of Precambrian land biota is extremely sparse). The postulated evolution of land biota corresponds to the best current inferences from fossils and molecular phylogenetic reconstructions (see earlier discussion) that put hyperthermophiles at the base of the tree of life. Thus, the earliest land biota likely included hyperthermophilic methano-gens, chemoautotrophs consuming atmospheric carbon dioxide in a pressure cooker atmosphere, along with traces of hydrogen produced by photodissociation of water. Atmospheric hydrogen levels of 1% for 108 years is apparently compatible with tropospheric photochemistry (Walker 1977). Other members ofthe earliest land biota probably included other hyperthermoph-iles such as the ancestors of filamentous Chloroflexus aurantiacus (a photoauto-troph feeding on hydrogen, hydrogen sulfide, and carbon dioxide), now found in hot springs, as well as Pyrodictum and Acidanus (consumers of hydrogen and free sulfur). Other living models of this earliest land biota are found in the diverse community of strict anaerobes living in hot springs (submarine hydrothermal environments), often in matlike colonies on rock substrates. The postulated ancient communities covered the early Archean continents and volcanic islands with similar microbial mats, which dried out between rain events, only to be revived when wet.

Cyanobacteria, along with oxygenic photosynthesis, likely joined this land consortia not later than 3.5 Ga, judging from the inferred fossil record of the Warrawoona cherts (Schopf 1992, 1993). By providing aerobic microenvironments, these mats likely were home for the earliest aerobic eucary-otes, although their emergence was likely delayed by some billion years because of ambient temperatures above their upper limit of metabolism (60°C). These late Archean/early Proterozoic communities were plausibly antilichens, reversing the contemporary lichen spatial relationship between pho-toautotrophs and heterotrophs. Modern procaryotic mats similarly contain diverse eucaryotes such as diatoms (Brown et al. 1985).

By the mid-Proterozoic, with the rise of atmospheric oxygen, thicker, more productive eucaryotic mats likely dominated the land biota, with stronger soil-stabilizing power and water retention. By the late Proterozoic, new organisms such as primitive lichens likely emerged, with primitive fungi such as chytrids constituting the host for cyanobacterial or eucaryotic cells (modern lichens have with few exceptions Ascomycetes or Basidiomycetes host fungi). Even these primitive lichens may have been predated by actino-lichens, having heterotrophic bacteria, actinobacteria, hosts; modern descendants have hyphae-like growth, forming matlike colonies. By the early Paleozoic, true plants emerged, first bryophytes.

The elevation of soil pCO2 by prevascular plant land biota—first microbial, then bryophytic—was perhaps even comparable with later vascular plant-dominated soils, thereby contributing to chemical weathering (Keller and Wood 1993; Yapp and Poths 1994). The appearance of vascular plants in the early Silurian, with their global expansion by Carboniferous times, may have raised the soil pCO2 even higher by root respiration and decay of organic matter. More important was the likely increase in soil thickness and stabilization relative to the earlier microbial and bryophyte land biota (Wright 1990). A new factor leading to an acceleration of chemical weathering and cooling arose, the rhizosphere (the association of mycorrhizae and plant root), with its multifold factors (organic acid and chelating agent production and increase in contact with soil minerals) (Berner 1995a; Retallack 1997). The emergence of a significant rhizosphere is documented from early Devonian fossils (Elick et al. 1998). Perhaps the great glaciation of the Carboniferous/Permian was related in part to this new factor (Algeo et al. 1995; Berner 1995a).

Much later, the emergence of angiosperms and their competition with conifers by early Tertiary may have further increased weathering rates, apparently because of the mycorrhizal relationships of angiosperms, which led to more efficient nutrient extraction from minerals, increasing the rate ofCO2 removal from the atmosphere (Volk 1989b). However, Robinson (1991) argued that the contemporary field data used by Volk to infer higher weathering rates for angiosperms was biased by more weatherable rocks in these watersheds. She concluded that there is no demonstrable difference between the chemical weathering rates of angiosperms and gymnosperms. Thus, it is problematic whether the increase of angiosperms contributed to a global cooling trend, which indeed is apparent in the last 100 million years. The later cooling may be more linked to a decrease in sea-floor generation rate (and outgassing). On the other hand, the emergence of widespread grasslands, with their deep root systems and sink for silica, could have promoted higher weathering rates and global cooling in the mid-Miocene ( Johansson 1993, 1995).

The preferred scenario for surface temperature as a function of geologic age is shown in figure 8-3. Second-order perturbations arising from such factors as continental drift, episodic burial of organic carbon, and pulses of intense volcanism are ignored (particularly relevant in the Phanerozoic, when atmospheric pCO2 levels decrease). Marked cooling in the mid-Proterozoic is indicated, consistent with the apparent emergence ofMetazoa at 1.0 to 1.5 Ga (Chapman 1992; Morris 1993) and paleotemperatures of 25 to 43°C on 1.1 to 1.2 Ga cherts (Kenny and Knauth 1992). This cooling may have been the result of the rise of atmospheric oxygen, as already pointed out, and more extensive colonization of the land surface by algal mats (Beeunas and Knauth 1985; Horodyski and Knauth 1994). This in crease in terrestrial biotic productivity, along with the onset of frost wedging in mountains, likely substantially increased the biotic enhancement of weathering. Note that although frost wedging is an abiotic process, here I argue that it commenced as an important physical weathering factor as a result of global biotically mediated cooling, thus bringing in this factor as a component of the cumulative global biotic enhancement of weathering. From the pattern of inferred variation of the U/Th ratio in the depleted mantle over geologic time, Collerson and Kamber (1999) suggest that the reinjection of oxidized uranium begins at about 2 Ga, accelerated by enhanced biotic weathering and continental erosion, consistent with the above argument.

Although the first definitive fossil evidence for Metazoa dates at 0.65 Ga (McMenamin and McMenamin 1990), an older problematic record does exist (Hofmann 1994; Robbins et al. 1985; Breyer et al. 1995; Fedonkin et al. 1994; Seilacher 1997; Seilacher et al. 1998). The story of the earliest Metazoa may be analogous to that of Eucarya (Sogin et al. 1989) in the unlikely preservation of microscopic soft-bodied organisms as fossils. Davidson et al. (1995) noted that "micrometazoan ancestors would not have left a fossil record because of their small and probable lack of skeletonization." These ancestors are proposed to "constitute a cryptic pre-Ediacaran record." The emergence of Metazoa at 1 Ga or earlier are supported by the results of a molecular phylogenetic study (Wray et al. 1996), although their conclusions were challenged by Ayala et al. (1998), who argued that the divergence of metazoan phyla occurred no earlier than 670 million years based on their interpretation of molecular clocks. However, even if they are right, their analysis apparently does not rule out an earlier primitive metazoan emergence prior to the divergence of phyla.

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