Silicon (2009) 1:65–77 DOI 10.1007/s12633-009-9014-2
REVIEW
The Potential use of Silicon Isotope Composition of Biogenic Silica as a Proxy for Environmental Change Melanie J. Leng & George E. A. Swann & Martin J. Hodson & Jonathan J. Tyler & Siddharth V. Patwardhan & Hilary J. Sloane
Received: 8 June 2009 / Accepted: 6 July 2009 / Published online: 18 July 2009 # Springer Science + Business Media B.V. 2009
Abstract Silicon isotope geochemistry is a relatively new branch of environmental change research. Here we review the recent developments in the preparation of materials, analytical methods and applications of stable silicon isotope geochemistry in the most common types of biogenic silica currently being analysed. These materials are: diatom, radiolarian and siliceous sponges in lake and ocean sediments and plant phytoliths which are preserved in soils. Despite analyses of Si isotopes being carried out on rocks and minerals since the 1950's and the increasingly widespread use of Si isotopes since the 1990's, to date only a relatively small number of studies have applied Si isotope ratios to environmental change. In lake and ocean sediments the analysis of Si isotope ratios from biogenic materials has the potential to provide an important source of palaeoenvironmental inforM. J. Leng (*) : G. E. A. Swann : H. J. Sloane NERC Isotope Geosciences Laboratory, British Geological Survey, Nottingham NG12 5GG, UK e-mail:
[email protected] M. J. Leng School of Geography, University of Nottingham, Nottingham NG7 2RD, UK M. J. Hodson School of Life Sciences, Oxford Brookes University, Headington, Oxford OX3 0BP, UK J. J. Tyler Department of Botany, The Natural History Museum, Cromwell Road, London SW7 5BD, UK S. V. Patwardhan School of Science and Technology, Nottingham Trent University, Nottingham NG11 8NS, UK
mation, especially where carbonates are not preserved. In plants and soils few studies have used Si isotopes, but important advances have recently been made in the understanding within plant fractionations. These may be useful in the application of Si isotopes in phytoliths to archaeological and palaeoenvironmental contexts. Keywords Si isotope ratios . Biogenic silica . Environmental change
1 Introduction Four isotopes of silicon exist in the natural environment: 28Si, 29 Si, 30Si (which are stable) and 32Si (which is radioactive). The relative abundances on Earth of the stable isotopes are 92.23% (28Si), 4.67% (29Si), and 3.10% (30Si) [1]. The silicon isotope ratios (30Si/28Si and 29Si/28Si) of materials are expressed on the delta-scale (δ) in terms of per mille (‰): d¼
Rsample =Rreference 1 :103
ð1Þ
Where R is 30Si/28Si and 29Si/28Si and ‘reference’ means the appropriate universally-accepted reference material. The ‘δ’ for each element takes its name from the heavy isotope, thus δ30Si. For silicon the standard reference material is NBS-28. Measurements of δ30Si (30Si/28Si) and δ29Si (29Si/28Si) are generally done simultaneously, and there is a mass dependent fractionation between the two ratios, so that δ30Si= 1.96×δ29Si, with measurement of both ratios being a good analytical indicator of sample purity. Mass spectrometry studies of silicon isotope variation in the natural environment started in the 1950’s [2–5]. Significant methodological advancements came as a result of the Apollo lunar landing program in the late 1960’s and
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70’s [6, 7] and research has continued since then looking at improvements in the methodology [8–10]. The isotope composition of a variety of materials including plant phytoliths, diatoms, siliceous sponges, dissolved silicon in waters as well as igneous rocks has now been determined [11–20]. Within the aqueous environment dissolved silicon is present as orthosilicic acid (e.g. H4SiO4), where it is utilised by a variety of organisms to produce biogenic silica (e.g. Fig. 1) as a structural element, including higher plants, which deposit phytoliths, unicellular phytoplankton (diatoms, chrysophytes, silicoflagellates), zooplankton (radiolaria) and sponges [21]. Often the availability and concentration of silicic acid and the production of biogenic silica is a critical factor in the survival of these organisms. The silicon cycle is one of our major biogeochemical cycles (Fig. 2, adapted from [22]), and is especially important in environmental change research because of its linkages with the carbon cycle. It has been estimated that the global annual fixation of phytolith silica (60–180 Tmol yr−1) is on
the same order of magnitude as the amount annually fixed in ocean diatom communities (240 Tmol yr−1), with soils containing orders of magnitude more amorphous Si, primarily buried as phytoliths [23]. In both marine and terrestrial environments, recycling of amorphous silica is the driving force behind the Si cycle. The majority of research on silicon isotopes and silicic acid utilization has been conducted on diatoms (δ30Sidiatom) (Fig. 1). During the uptake and conversion of silicic acid into particulate hydrous silica (biomineralization), the lighter 28Si is preferentially incorporated into the silica over the heavier 29Si and 30Si. Therefore, within the photic zone of a water body, progressive silica uptake leads to an increase in the δ30Si of both the dissolved and particulate phases. Thus δ30Sidiatom can be related to nutrient utilisation, which is largely driven by both nutrient supply and demand. However, mixing between individual water masses containing different concentrations and isotope composition of dissolved silicon (δ30SiDSi) [24], as well as the impact of changes in iron availability, which can
Fig. 1 Forms of biogenic silica: Scanning Electron Microscope images of centric diatoms with tephra (contaminant) from Lake Tilo in Ethiopia (a), a well preserved diatom frustule (Genus Coscinodiscus) from the North Pacific Ocean (b), Transmission Electron Microscope images of Aulacoseira granulata var angustissima from a freshwater diatom monoculture (c) and (d), Light micrographs illustrating the two main
types of Si deposition in plants—lumen and wall: (e) shows the ashed remains of a wheat culm (stem), (f) shows a single dendriform phytolith isolated from the wheat inflorescence. All the epidermal cell walls are silicified, including the long cells, the short cells in between them, and the single stoma
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Fig. 2 The global silicon cycle. The size of the boxes are not to scale and do not represent the relative size of each silicon pool
influence biogenic silicon demand [25, 26], may complicate any interpretation. Similar effects are also likely in lacustrine environments, but on a local scale, with variations in catchment geology, river input, weathering, water residence and timing of seasonal diatom blooms all affecting δ30SiDSi and δ30Sidiatom. To date though, few studies have investigated δ30Sidiatom in lakes [27, 28]. In higher (vascular) plants silicon is taken up by plant roots from the soil in the form of monosilicic acid, which is then transported through the plant in the transpiration stream [29]. Phytoliths are formed in plant cells when the concentration of soluble silicon increases above about 2 mM, and solid amorphous silica is deposited. Thus phytoliths take on the size and shape of the compartment in which they are deposited (e.g. Fig. 1e,f), varying between cell types, different organs within a plant and between different plant species [30]. Plant species vary in the amount of silicon that they accumulate [31], and consequently in their phytolith production. Phytoliths have been much used in plant taxonomy [32], archaeology [33] and palaeoecology [34], and until recently most of this work has relied on determining morphological differences between phytoliths from different taxa. There has, however, been increased interest in using the chemical features of phytoliths as markers, and the use of silicon isotopes should be seen in that context [17]. Only in the last five years have
we begun to understand the processes that are involved in silicon isotope uptake and distribution within a plant, and as yet there have been no attempts to use this knowledge in archaeological or palaeoecological work. Here we will review the use of silicon isotopes in environmental change research, considering the methodologies employed and examples from the work that has already been undertaken.
2 Chemistry, Structure, and Biosilicification of Biogenic Silica The chemical nature of biogenic silica can be represented as Si(OSi≡)n (OH)4−n where n 50 μm) opens the possibility that it may become viable to analyse individual diatom frustules or phytoliths. For this to be viable, however, it becomes necessary to reduce the analytical errors associated with SIMS which are currently higher than those of other techniques [93].
5 Applications of Silicon Isotopes to Environmental Research 5.1 Catchment/lakes In the oceans there have been studies that have confirmed the link between diatom silicon utilization and Si isotopic compositions [98–100]. Consequently, silicon isotope data from the sedimentary record has been used as a proxy to reconstruct palaeoproductivity (or more strictly marine silicic acid use by diatoms relative to initial dissolved silicic acid concentrations, ie. [101, 102]. In continental waters, there have been relatively fewer silicon isotope investigations, and those reported tend to be on major rivers [94, 98]. Si in rivers comprises both dissolved and particulate matter, and measurement of both allows an assessment of weathering as well as productivity related fractionation. Other studies have focused on the relationship between climate, diatom productivity, and lake mixing regimes [27]. These studies, whilst highlighting the complexity of lake systems (potentially having several weathering component fluxes) show that silicon isotope ratios are consistent with Si concentrations [98]. They also suggest that Si isotope fractionation is independent of species and temperature, offering potential information on changes in nutrient supply and limnological change. There are currently rather few studies on silicon isotopes from biogenic components of lake sediments. One of the first is from Lake Rutundu, Mt. Kenya, over the last ~38 kiloannum (ka) which attempted to use a small, well studied lake/ catchment in order to highlight processes that may have operated at biome to continental scales during the Late Quaternary [28]. A combination of lake-sediment fluxes and stable-isotope (δ13C, δ15N, δ18O, as well as δ30Si) data, were used, and the authors showed that under glacial conditions (38.3–14.3 ka), high diatom productivity was maintained by substantial transport of dissolved SiO2 and soil nutrients from a sparse, leaky, terrestrial ecosystem. During the following period of high monsoon rainfall and seasonality (14.3 to 9.5 ka), rapid Si cycling by fire-prone grassland was associated with substantial aeolian transport of opal phytoliths
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by smoke plumes, but greatly reduced nutrient flux in runoff. Invasion of tall, subalpine shrubs after 9.5 ka further enhanced landscape stability, leading to very low sediment fluxes of both phytoliths and diatoms. Given that there is much discussion in the public arena regarding silicon isotopes in palaeolimnology, we are likely to see a surge of interest in the years to come especially since analysis of oxygen isotopes in biogenic silica has become so popular [87]. 5.2 Oceans Relative to other systems, the majority of biogenic silicon isotope studies over the past decade have centred on the marine environment. Early measurements of δ30Si were employed as a tracer of modern day biogenic silica production and dissolution within the water column [103]. Whilst this has been largely replaced by the use of the radioactive 32Si (half-life= c. 140 years) [104], δ30Si remains essential when investigating changes beyond the last few hundred years. Understanding the controls on the isotopic composition of δ30SiDSi used during biomineralisation is essential for interpreting spatial and temporal changes in the silicon isotope composition of biogenic silica in the marine environment. As well as the marine silicon cycle itself, this necessitates a holistic view of the other components of the global silicon cycle which fundamentally control the supply of silicon to the oceans, e.g., the biosphere, terrestrial, riverine and lacustrine systems [23]. For example, variations in δ30Si of biogenic silica can arise due to the mixing of individual water masses containing different values of δ30SiDSi, in particular the mixing of surface water and deep water [24]. Contemporary studies have demonstrated a relatively large range of marine δ30SiDSi values from +0.5‰ to +3.2‰ [20, 24, 98–100], whilst offsets of up to 0.4‰ have been documented between Atlantic and Pacific deep waters over recent glacial-interglacial cycles [98, 101]. Other studies have focused on using ocean models to understand the global distribution and changes in marine δ30SiDSi [105, 106]. For example, using a simple modelling approach De La Rocha and Bickle [106] successfully demonstrated that whole-ocean changes in δ30SiDSi arising from variations in riverine input are within analytical error over the last 100,000 years. Since silicic acid is predominantly delivered to the oceans via riverine inputs, on the basis of these results δ30Si in biogenic silica can be used to monitor the marine silicon cycle over both the modern day as well as over glacial-interglacial cycles [106]. With regards to other forms of biogenic silica, a limited amount of work has been carried out on siliceous sponges [11, 107] and radiolaria ([67] and references within). Results from siliceous sponges indicate their potential for monitoring deep water changes in δ30SiDSi [11]. However, it remains unclear
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whether the δ30Si record from these organisms is subject to some form of inter-species vital effect [11]. By a significant margin, most marine research to date has focused around diatoms. Both laboratory and in field experiments have shown that measurements of δ30Sidiatom are subject to an enrichment factor, independent of temperature, pCO2 or other inter-species vital effects, of −1.1‰ to −1.9‰ [98, 99, 108– 111]. The information provided by δ30Sidiatom on nutrient utilisation and export production therefore provides an important link in understanding the role of the siliceous biological pump in transferring carbon into the deep ocean and regulating atmospheric concentrations of CO2 over glacial-interglacial cycles [12, 102, 112] as well as changes in the global silicon cycle over geological timescales [113]. Measurements from the Southern Ocean, for example, indicate that δ30Sidiatom were c. 1.0‰ lower in the Southern Ocean during glacial periods, in line with reduced diatom productivity, over the last three glacial-interglacial cycles [101, 102]. These records suggest a relative decrease in diatom silicic acid utilisation during glacials, possibly resulting in the build up of a concentrated pool of silicic acid across the Southern Ocean. The subsequent northward migration of this silicic acid during glacials would have enabled diatoms to dominate at lower latitudes during glacials than today, potentially increasing the net drawdown of CO2 into the ocean and lowering atmospheric pCO2 [102, 112]. However, whilst evidence exists to support such a series of events in the sub-Antarctic sector of the Indian Ocean, similar changes in δ30Sidiatom are not apparent in the sub-tropics making it unclear whether any pool of silicic acid extended into the sub-tropical region during glacial periods [114, 115]. 5.3 Plants Silicon isotopes are taken up by plant roots, and it is clear from work on rice grown in hydroponic cultures that considerable fractionation occurs as silicon is taken into the root. Sun et al. [116] report δ30Si of +0.3‰ in a culture solution, and −0.7‰ in the root of rice. A similar fractionation has been observed in banana [117]. This would imply that lighter isotopes enter the plant preferentially to the heavy isotopes. It has been suggested that this indicates a biologically mediated fractionation [116]. Recently, the genes encoding silicon transporters in the roots of rice, barley and maize [62–64, 66] have been isolated. It is possible that these transporters are responsible for the observed fractionation, maybe by discriminating against heavy isotopes, but this has yet to be determined. The relative accumulation of isotopes by plants depends to a large extent on the substrate on which they grow (Table 3). It is now clear that isotopic composition of phytoliths will be affected by factors such as the soluble silicon concentration of the medium [117], the weathering of the substrate [118], and
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Table 3 Published δ30Si and δ29Si from plant species illustrating variation due to growth in different environments, and fractionation between plant organs References
Plant Species
Environment
δ29Si from plants growing in one environment
Single and range of δ30Si from plants growing in one or different environments
Range of δ29Si between plant organs
Range of δ30Si between plant organs
[17]
Bread Wheat
Soil, UK
+0.44
+0.68
Between −0.62 and +1.18
Between −1.4 and +2.14
[119] [119] [119] [119] [119] [119] [119]
Soil, Soil, Soil, Soil, Soil, Soil, Soil,
Sweden Sweden Sweden Sweden Sweden Sweden Sweden
+0.12 +0.13 −0.07 −0.01 −0.15 −0.05 −0.05
+0.2 +0.22 −0.22 −0.1 −0.31 −0.18 −0.15
Soil, Sweden
−0.34
+0.56
[119] [119] [18]
Norway Spruce European Larch Downy Birch Rowan Goat Willow Lingonberry Lyme Grass (spring) Lyme Grass (autumn) Step Moss Lichen Bamboo
Soil, Sweden Soil, Sweden Seven soils, China
−0.08 −0.08
[120]
Rice
Four paddies, China
[107]
Horsetail
Three soil, California
[107] [118]
Bamboo Banana
[121]
Banana
Soil, California Two clay soils, West Africa Soil, Cameroon
−0.24 −0.14 Between −2.3 and +1.8 Between −0.5 and +0.4 (stems and leaves) Between +0.8 and +2.8 −1.4 Between +0.02 and +0.54
[119]
[121] [116]
Hydroponic culture Rice
Hydroponic culture
soil organic matter [19]. As yet there have been no attempts to relate the isotopic composition of phytoliths in plants to those isolated from the soil, and experiments in modern situations need to be undertaken. However, there is potential, in archaeological contexts, to use the isotopic composition of phytoliths to determine whether plants were grown locally or were imported from elsewhere. In order to assess this potential, experiments are required whereby plants are grown under controlled conditions which simulate past environments (soil, irrigation, temperature etc.), with subsequent comparison of fossil and contemporary phytolith isotope signatures. Recently, it has become clear that there is a large fractionation of silicon isotopes within plants (Table 3), and that the degree of fractionation appears to depend on the extent of Si accumulation. For example, fractionation is far greater in rice [116, 120], a heavy silicon accumulator, than banana [117, 121], a moderate accumulator. Lighter isotopes are preferentially deposited in phytoliths lower
Between −2.42 and −1.82
Between +1.0 and +3.3 Between −0.16 and +2.98
Between −0.18 and +1.09 Between +0.13 and +0.49 Between −0.7 and −0.44
down the plant, and a greater proportion of heavier isotopes are found within the xylem. This means that the heavier isotopes tend to accumulate towards the end of the transpiration stream. It has been suggested that Rayleigh fractionation is responsible for this phenomenon [120]. Again the question arises as to how this finding might be used by archaeologists and palaeoecologists. We would nearly always expect leaves to have different signatures to culms (stems) in grasses and cereals, but light microscopy will easily distinguish these phytolith types. This suggests that provenance of phytoliths might be the only use of silicon isotopes in phytoliths for archaeological research.
6 Summary Measurements of silicon isotope ratios from biogenic silica provide a potentially important source of environmental
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change information, which can extend our understanding of palaeolimnology and palaeoceanography in both high latitude regions and other sites containing biogenic silica where more traditional materials analysed for isotopes are absent. Recent years have been marked by several advances with regards to method development and our understanding of the silicon isotope signal. However, a number of other areas need to be addressed in order to improve the accuracy of silicon isotope based reconstructions. There is still a need to ensure purified materials are analysed since all silicate contaminants will contain Si, and there is also a need to further develop existing/new analytical procedures for analysing silicon which do not require fluorine reagents (i.e. MC-ICP-MS). Such steps will likely lead to increased numbers of laboratories developing analytical facilities and more research being undertaken, which will ultimately lead to an advancement in the science. At present, despite the considerable potential in using records of silicon isotope ratios to provide further insights into the nature of climatic/ oceanographic changes in carbonate free/biogenic silica rich regions, only a limited number of studies have been carried out on marine and lake sediments in comparison to other techniques [122]. To date, the rapid expansion of δ18Odiatom, particularly in palaeolimnology [87] has not been matched in measurement of silicon isotopes. This, however, may be readdressed in the near future by the development of a combined methodology for analysing δ18Odiatom and δ30Sidiatom on the same sample [10]. In recent years, rapid progress has been made on research within silicon isotope variation in modern plants. It is now evident that the substrate the plants are growing in will influence the isotopic composition of phytoliths. It is also firmly established that considerable fractionation occurs within plants, particularly in heavy silicon accumulators. However, much basic work remains to be done with modern plants to elucidate the mechanisms behind these phenomena. At present we do not even know if environmental factors such as water availability to plants will influence the Si isotopic composition of phytoliths. We would caution against rushing to determine silicon isotopic composition from phytoliths in soils and sediments from past environments until more is known about the processes that determine fractionation within the plant. Acknowledgements Dr. O. Deschaume is thanked for his help producing Figure 3. The authors wish to thank the four anonymous reviewers for their helpful comments.
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