The oxygen isotopic composition of phytoliths from tropical rainforest soils ( Queensland , Australia ) : application of a new paleoenvironmental tool

Introduction Conclusions References

Here, we explore the efficiency and limitations of using the oxygen isotopic composition of wood phytoliths (δ 18 O wood phytolith ) from tropical rainforest soils as a suitable proxy for atmospheric temperature and δ 18 O precipitation values, under conditions that are assumed to be non-evaporative.Introduction

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Full Phytoliths are micrometric particles (<60-100 µm of diameter) of amorphous silica that form within a matter of hours to days (Perry et al., 1987) inside or between the cells of higher plant tissues throughout the life of a plant.With plant decay, phytoliths are either incorporated into soils or exported to sediments via regional watersheds.Well preserved in oxidizing environments, phytoliths are recovered from modern to Tertiary materials (e.g.Stromberg et al., 2005).Through the morphological identification of phytolith types, modern and fossil phytolith assemblages have been shown to be a reliable tool for characterizing either the variety of inter-tropical grasslands and for inferring tree cover density at low elevation tropical sites (e.g.Alexandre et al., 1998;Boyd et al., 2005;Bremond et al., 2005aBremond et al., , b, 2008a, b;, b;Piperno, 2006;Lentfer and Torrence, 2007;Neuman et al., 2009).
In parallel with the increasing use of morphological assemblages of phytoliths as vegetation proxies, pioneering studies of the δ 18 O signature of phytoliths were conducted using harvested grasses grown under measured conditions of temperature, relative humidity, and isotopic compositions of plant and soil waters (Shahack-Gross et al., 1996;Webb and Longstaffe, 2000, 2002, 2003, 2006).Such studies have demonstrated, in agreement with previous results (e.g.Ometto et al., 2005), that the absorption of water by roots is a non-fractionating process and that in non-transpiring grass stem tissues, the equilibrium fractionation between water and phytolith is temperature-dependent; the temperature coefficient being of −0.33 ‰/ • C (recalculated from Shahack-Gross et al., 1996).They additionally evidenced that in transpiring tissues, 18 O enrichment of sap water increases with the inverse of relative humidity, which prevents the use of the δ 18 O signature of phytoliths from grass leaves as a function of temperature and the isotopic composition of soil water (δ 18 O soilwater ) (Webb and Longstaffe, 2000;2002;2006).These calibration studies, although extremely useful for our understanding of the temperature and soil-water δ 18 O signals carried by phytoliths, have not lead to paleoenvironmental reconstructions due to the fact that phytoliths from both transpiring and non-transpiring grass tissues are not morphologically distinguishable.Introduction

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Full In light of these results, we measured the oxygen isotopic composition of rainforests phytolith assemblages collected from the soil humic horizon (or soil top) in the Wet Tropics of North-East Queensland, Australia.We investigated whether a temperaturedependant relationship between the oxygen isotopic composition of wood phytoliths (δ 18 O wood phytolith ) and δ 18 O precipitation values can be obtained and if it is close to the one previously obtained between isotopic compositions of harvested grass phytolith and plant water.
We focused on Queensland rainforest soil top phytolith assemblages with the overall objectives of paleoenvironmental reconstructions for the following reasons: (1) in small lake's catchments covered by rainforest such as the ones found in Queensland (e.g.Haberle, 2005;Kershaw et al., 2007), phytoliths from eroded soil tops are expected to constitute the main source of fossil phytoliths in sediments; (2) the weak concentration of phytoliths in rainforest soil tops (a few ‰ in weight, Alexandre et al., 1997) and in the suspended load of tropical rivers (less than 5 ‰ in weight, Carry et al., 2005) suggests that the mean phytolith concentration in lake sediments is also on the order of a few ‰ in weight.Taking into account the amount of phytoliths required for δ 18 O analyses (several mg) and using the lake sediment accumulation rates commonly observed in Queensland rainforest environments (cm 100 yr −1 ; e.g.Haberle, 2005;Rieser and Wust, 2010), the fossil phytolith assemblages provided using sampling intervals of several centimeters in lake sediment cores should encompass 100s of years.For comparison, mean age of 100s of years was measured for bulk organic matter (OM) from one of the soil top samples investigated in the present study ( 14 C mean age of sample #18 was 610 ± 15 yr BP; AMS-14 C UCIAMS #75077), and can be reasonably expected for phytoliths (Alexandre et al., 1999); direct 14 C AMS-measurements of phytolith-occluded carbon, although very promising, being still in a calibration phase (Prior et al., 2005;Rieser et al., 2007;review in Santos et al., 2010); (3) phytolith assemblages from rainforests should carry a δ 18 O signature suitable for paleoenvironmental interpretations.
Indeed, they are constituted by more than 80 % of a type of phytolith (Alexandre et al., 1997;Runge, 1999;Bremond et al., 2005b), named as Globular granulate after the Introduction

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Full international code for phytolith nomenclature (Madella et al., 2005), mainly formed in the xylem of the wood of rainforest trees (Scurfield et al., 1974;Kondo et al., 1994) (Fig. 1).Wood is a non-transpiring organ and the isotopic composition of xylem water transported from the roots throughout the plant is not modified by transpiration.Hence, the δ 18 O wood phytolith signature should relate to the δ 18 O soil water signature; (4) the area of the wet tropical rainforests of northeast Queensland, confined to 18 000 km 2 , bordered by the Pacific to the east and dry woodlands and savannas to the west, encompasses a large altitude gradient (0-1600 m a.s.l.) with steep gradients in temperature and precipitation (16-25 • C and 1000-8000 mm yr −1 ) over a short distance (DASETT, 1986).Such gradients are ideal for assessing climatic and altitude controls on the δ 18 O signature of rainforest phytolith assemblages.

Geomorphic features
The Wet Tropical rainforests of Queensland extend from 16 • 27 to 17 • 38 S and from 145 • 19 to 145 • 51 W and have been protected as a world heritage area since 1988.
The main distribution of rainforests in northern Queensland straddles three geologic and geomorphic regions that are NW-SE or N-S oriented (Fig. 2a

Climate features
The region is located at the southern limit of an area that is influenced by the Australian Summer Monsoon (ASM).The climate pattern is controlled by the position of the Intertropical Convergence zone (ITCZ) and the monsoon circulation.Over 80 % of mean annual precipitation falls during the November-March interval, supplied by NW monsoonal winds during the earlier phase of the ASM, by S-E trade winds during the later phase of the ASM (Godfred-Spenning and Reason, 2002) and by occasional cyclones.In the NW-SE oriented eastern highlands, there is a very pronounced orographic influence with higher precipitation averages centred around the highest peaks and their eastern slopes, windward to the moist prevailing SE airstream.Mean annual precipitation reaches 4000 mm yr −1 along the coast, increasing with elevation to reach 11 000 mm yr −1 at Mt Bartle Frere, and falls dramatically inland (1200 mm yr −1 on the western Tableland) (Fig. 2c).At higher ranges (from 1000 m a.s.l.), fog and orographic cloud layers shroud the summits of the mountains and maintain moist conditions throughout the year.Fog deposition accounts for approximately 40 % of the water reaching the forest floor (Hutley et al., 1997).The amount of water derived from the canopy interception, also called "cloud stripping" by these high-altitude "cloud-forests", can account for 66 % of precipitation during the dry season (McJannet et al., 2007a, b).Coastal humidity averages 78 % in the summer but often reaches values higher than 90 %.Mean annual temperature (MAT) exceeds 24 • C along the coast and falls to below 21 • C on the tableland, and to below 17 • C in the highest ranges (DASETT, 1987;Moss and Kershaw, 2000;Godfred-Spenning and Reason, 2002;Robertson et al., 2005).

Vegetation features
Rainforests of the wet tropics of Queensland are classified into 27 structural categories that contain more than 3000 plant species from 210 families (Tracey, 1982;Webb and Tracey, 1994).Soils were sampled from areas that support seven rainforest categories Introduction

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Full (Table 1), numbered and described by Tracey (1982)  Full

Materials
Seventeen samples were collected from poorly drained podzolic soils, along four altitudinal transects from 70 to 1283 m a.s.l.(Table 1; Figs. 2 and 3).The sampling method consisted of collecting individual sub-samples of the upper 2 cm of the soil humic horizon (litter excluded), at random intervals, over an area of 5 × 5 m.Sub-samples were mixed together.Three transects where samples were collected are located in the eastern highlands, east Atherton, while the fourth transect (Northern Cairns) is located on the coastal belt.The "Bartle Frere" transect is located on the North-Western slope of Bartle Frere South Peak (1615 m a.s.l.), above Russell River oriented NW-SE when its upstream section crosses the upper rainforest area.All of the sampled points are leeward relative to the dominant S-E trade winds.The "Palmerston Highway" transect is located in the eastern section of the eastern highlands, above the Johnston River valley (NW-SE).The sites sampled are windward to the dominant S-E trade winds.The "Mt.Edith" transect is located on the southern slope of Mont Edith (1149 m a.s.l.).The sites sampled are windward, but the S-E trades first pass over the Bellenden Ker range (maximum altitude of 1545 m a.s.l., mean altitude of 1100 m a.s.l.) before reaching the tablelands and flowing up to Mt. Edith.The "Northern Cairns" transect is located on the coastal belt.The sampled points are windward (Fig. 3) relative to the dominant S-E trade winds.
Long term climate means at the sampling sites (Table 1) were obtained from regional, digital maps of bioclimatic variables that were created using the ANUCLIM software (McMahon et al., 1995) which uses a Digital Elevation Model (DEM) and meteorological data from a large number of stations over variable time periods (from several to many decades until approximately 2000) to estimate the climate variables for each grid cell in the DEM (0.1 km 2 ).Introduction

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Full

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Full cuneiform bulliform cells, and short cells types; and (4) types without taxonomic significance (unclassified).Abundances of the classified phytolith categories were expressed as a percentage of classified phytoliths, while the abundance of unclassified types was expressed as a percentage of the sum of counted phytoliths (Table 2).Two to 60 µm size particles of quartz, iron oxide and charcoal, as well as larger thin remains of OM, were sometimes recovered and hence counted (Table 2).

δ 18 O silica measurements
Phytoliths are hydrous silica particles that contain exchangeable oxygen mostly in hydroxyl groups (Labeyrie and Juillet, 1982;Perry and Keeling-Tucker, 2000).In order to evaluate the amount of exchangeable oxygen, to fix the isotopic composition, and to calculate the isotopic composition of non-exchangeable oxygen (δ 18 O silica ) a controlled isotopic exchange procedure (CIE) was carried out.Two aliquots of 1.6 mg of the same sample were exchanged with vapor from two waters of a known isotopic composition (Crespin et al., 2008).Oxygen extractions were then performed using the IR Laser-Heating Fluorination Technique as described in Alexandre et al. (2006) and Crespin et al. (2008).Oxygen gas samples were directly sent to and analyzed by a dual-inlet mass spectrometer (ThermoQuest Finnigan Delta Plus).The oxygen isotopic results are expressed in the standard δ-notation relative to V-SMOW.The measured δ 18 O values of each sample (δ 18 O measured1 , δ 18 O measured2 ) were corrected on a daily basis using a quartz lab standard (δ 18 O Boulang é 50−100 µm = 16.36 ± 0.09 ‰).One to three aliquots of a phytolith lab standard, MSG 40 (Crespin et al., 2008), were run per measurement session in order to verify the effectiveness of the CIE procedure.Replicate analyses of the samples yielded a reproducibility for δ 18 O measured better than ±0.5 ‰, except for one sample (#19) (Table 2).The uncertainty or limit of the 90 % confidence interval associated with the calculation of δ 18 O silica using a Monte-Carlo simulation, was ±0.5 ‰ (Crespin et al., 2008).During the calibration period, replicate analyses of the international standard NBS 28 gave an average of 9.6 ± 0.17 ‰ (1σ, n = 13).Introduction

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Full A recent inter-laboratory comparison for oxygen isotopic composition of hydrous biogenic silica has evidenced that, when the CEREGE CIE is performed, a methodological bias occurs leading to abnormally high fractionation between the vapour and the exchanged oxygen (∆ vapour−Oexchanged ) (Chapligin et al., 2011).Several tests were conducted to examine possible impacts of water-vapour exchange conditions, the amount of silica subjected to exchange and the rate of silica dehydration.None of these tests produced significant changes in the δ measured values.Additionally, vapour-silica exchanges were carried out at three temperatures ( 140• C, 200 The consistency of this correction was verified using an independent data set previously obtained at CEREGE for fresh water diatoms from Annecy Lake (Crespin et al., 2010).The obtained relationship between ∆ diatoms−lake water and temperature of the lake allowed defining a corrected relationship: In a ∆ diatoms−lake water vs. t diagram, the obtained corrected fractionation line is shifted towards lower values of ∆ diatoms−lake water , lower but close to fractionation lines previously obtained for fresh water diatoms (e.g.Brandriss et al., 1998;Moschen et al., 2005;Dodd and Sharp, 2010).

Corrections on δ 18 O silica for obtaining δ 18 O wood phytolith
Since some rainforest phytolith assemblages contained a small amount of 2-60 µm particles that were not wood phytoliths, corrections were made to calculate δ 18 O wood phytolith values from δ 18 O silica values.
Correction for the presence of quartz particles: Weight correction was made for the amount of quartz particles (less or equal to 3 % of the counted particles in 7 samples) given the lowest δ 18 O quartz value of 8 ‰ attributed to the detrital quartz of metamorphic origin (e.g.Alexandre et al., 2006), and the respective densities of quartz (2.6) and phytoliths (2.3) (Table 3).
Correction for the presence of grass phytoliths: Phytoliths from the grass understorey may originate from transpiring grass leaves and may have δ 18 O values slightly enriched relative to wood phytoliths.Webb and Longstaffe (2002) previously demonstrated for a grass species collected from across the North American prairies that the 18 O enrichment of leaf phytoliths relative to stem phytoliths increases with the inverse of relative humidity: According to (Eq.4), for the 0.6-1 range of relative humidity calculated by the ANU-CLIM software for the sampled sites (Table 1), leaf phytolith 18 O enrichment should range from 5 to 0.5 ‰.In the absence of measured data on relative humidity in the under-storey, values from Table 1 were used in (Eq.4) to correct the obtained δ 18 O silica values for the presence of grass phytoliths.This correction is expected to be maximal.Indeed, in the lower canopy vapour pressure deficit is low (0-0.Amazonian rainforest (Ometto et al., 2005).Results with and without grass phytoliths corrections, were discussed.
Correction for the presence of palm phytoliths: To our knowledge, δ 18 O water signatures in forest palm stems and leaves have never been measured.If 18 O enrichment is similar for palm and grass phytoliths, (Eq.4) should also be used to correct the obtained δ 18 O silica values for the presence of palm phytoliths.Results with and without palm phytoliths corrections, were discussed.
No correction for the presence of unclassified phytoliths and charcoal: Unclassified phytoliths were assumed to mainly originate from tree wood and were not corrected for.The occurrence of charcoal particles was not corrected either since charcoal is largely made of carbon and should not contribute to the oxygen yield.
No correction for the presence of organic matter (OM): Thin organic matter remains, most of them with surface ranging from 100 to 200 µm, were counted.They accounted for less or equal to 5 % in 5 samples and for 12 % in one sample (Table 2).No correction was made for their low proportion as they should be oxidized at 1100 • C during CIE sintering step, and removed from the analyzed sample.

Estimation of atmospheric temperature and δ 18 O precipitation values
Ideally, matching the time span recorded by the soil phytolith assemblages would require us to obtain, for each of the sampled sites, measurements for 100s of years for MAT, δ 18 O soil water and/or δ 18 O precipitation values, which is unrealistic.Therefore, estimates were made, as justified below.
Regarding atmospheric temperature, measurements back to 1910 revealed a fast increase of MAT of +0.1 • C/10 y, mostly during the second half of the 19th century in Queensland (Suppiah et al., 2001).To our knowledge there is no other continuous record of temperature from any tropical Australian site.A dendroclimatological study showed evidence of a much slower increasing trend of +1.5 • C since the 16th century in New Zealand (Cook et al., 2000).From these records, the assumption was made that modern long term MAT values extracted from the GIS database (Table 1) should Introduction

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Full only slightly overestimate (by less than 1-2 • C) the mean value for the last 100s of years; this in a similar way for all the sampled sites.
Regarding precipitation, tree ring reconstruction from the Atherton tableland (Queensland) revealed no long trend since 1861 (Heinrich et al., 2008).In the absence of δ 18 O precipitation and/or δ 18 O soilwater measurements from Queensland rainforests, assumption was made that the long term depth-weighted δ 18 O soilwater values average the long term amount-weight mean annual δ 18 O precipitation values.Indeed, because Australian rainforests are characterized by low radiation levels due to frequent occurrence of fog and low clouds and a low range of the vapor pressure deficit for much of the time, soil evaporation and understorey evaporation are expected to be low (Hutley et al., 1997).For comparison, δ 18 O soil water values measured in the Amazonian rainforest were shown to be close to δ 18 O precipitation values and to lie along the local isotopic meteoric water line (Girard et al., 2000).To infer modern δ 18 O precipitation values, direct δ 18 O measurements could have been performed from non-evaporative surface water (e.g.river, spring water) (Lachniet and Patterson, 2009).However, given the confined sampled area, surface waters did not show sufficient variations in distance from the water source to record rapid changes in δ 18 O precipitation values with elevation.Finally, we calculated for each of the sampled site the amount-weighted mean annual δ 18 O precipitation value from latitude (LAT), and altitude (ALT) according to the relationships established by Bowen and Wilkinson (2002) from long term data obtained from the International Atomic Energy Agency-World Meteorological Organization Global Network for Isotopes in Precipitation (GNIP) database (IAEA/WMO, 1998).Driving factors, such as the origin of the air masses, the local surface air temperature, the latitude, the "distance from coast", and amount effects (e.g.Dansgaard, 1964;Gat and Matsui, 1991;Rozanski et al., 1993;Fricke and O'Neil, 1999;Gat, 2000;Gat et al., 2001;Poage and Chamberlain, 2001) are implicitly taken into account in these relationships.For stations located <200 m a.s.l.(Eq.5) and >200 m a.s.l.(Eq.6) the relationships are respectively expressed as: Introduction

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Full (Eq.5) and (Eq.6) were slightly modified to take regional conditions into consideration: (1) in humid tropics, δ 18 O values of precipitation from air masses lifted over high mountains appear to be controlled by the cumulative rainout upwind of collecting stations (Lachniet and Patterson, 2009).The dominance of this rainout process would result, for the windward stations, of an altitude effect of −0.002 ‰/m a.s.l. as measured worldwide (Gonfiantini et al., 2001;Lachniet and Patterson, 2002;Siegenthaler and Oeshger, 1980;Bowen and Wilkinson, 2002;Lachniet and Patterson, 2009), and for the leeward stations, of the delivery of moisture with low δ 18 O values (Lachniet and Patterson, 2009), also called "shadow effect"-δ 18 O precipitation values are lower than those predicted by altitude effects alone (Rietti-Shati et al., 2000;Longinelli et al., 2006).In order to take into account such a cumulative rainout, especially at leeward sites, the cumulative difference of the altitude crossed by the dominant S-E trade winds before reaching the sampled sites (|∆ALT|) was used for estimating δ 18 O precipitation both at the windward and leeward points (Fig. 3; Table 1).The procedure resulted in an inversed isotopic vertical gradient for the leeward slope of Mt.Bartle Frere, neglecting the role of increasing temperature as elevation decreases.This may have led to an underestimate for δ 18 O precipitation values especially at leeward low elevation sites; (2) the few isotopic studies investigating the water cycle in Amazonia, Costa-Rica, or Cameroon rainforest areas found an inland 18 O gradient of −0.08 ‰/100 km (for areas less than 230 km inland) which reflects the influence of recycled continental moisture (mainly from evapo-transpiration) on the isotopic composition of rain events (Salati et al., 1979;Gat and Matsui, 1991;Martinelli et al., 1996;Njitchoua et al., 1999;Lachniet and Patterson, 2002).Cloud striping that occurs in Queensland rainforests and is not taken into account in the inland gradient should additionally contribute to the rainwash that reaches the soil zone.Water trapped on leaves and undergoing evaporation would contribute to 18 O enrichment of soil water and to 18 O depletion of the remaining Introduction

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Full generated vapor that would, in turn, lead to 18 O depletion of the following rainfall event, finally increasing the inland 18 O gradient.In light of the above considerations, the inland 18 O gradient term (−0.08 (DIST)) was added to (Eqs. 5 and 6), keeping in mind that it could be underestimated for cloud forest sites.
Finally, for the <200 m a.s.l. and >200 m a.s.l.sites, the calculation of δ 18 O precipitation values was as follows: • 25 E; 594 m a.s.l.) located in the Atherton tablelands (Fig. 2).We assumed that weighted mean seasonal values were close to the weighted mean annual values, since most of the precipitation occurs during the rainy season.We checked that the δ
Quartz, OM, Fe oxides and charcoal particles were present in low proportions in few samples (Table 2).additionally corrected for the presence of grass and palm phytoliths (Table 3).When values from the leeward transect of Bartle Frere are excluded a modest correlation (p-value < 0.05) appears (Fig. 5):

Discussion
Whether corrections for the presence of quartz particles, grass and palm phytoliths are made or not on δ 18 O silica values, ∆ 18 O wood phytolith−precipitation values obtained for the windward transects showed close negative linear relationships with MAT expressed by their mean temperature coefficients ranging from −0.4 to −0.5 (±0.2) ‰/ • C and their y-intercept ranging from 46 to 48 (±4) ‰.These similarities showed the weak influence of the corrections for non-wood phytoliths on δ 18 O values of rainforest phytolith assemblages.

The obtained temperature-dependant relationship: comparison with previous studies
Equation ( 12) can be compared with the temperature-dependant relationships previously published for different silica-water couples (Fig. 6).
If we focus on the entire temperature-dependant relationships (fractionation lines) Fig. 6 recalls that there is no uniform relationship for the different silica-water couples.Several factors were previously accounted for these discrepancies: diatom frustules that show higher dissolution rate than that of phytoliths (Fraysse et al., 2009) may have been subject to early diagenesis which could have impacted their isotopic composition after deposition (Schmidt et al., 2001;Dodd and Sharp, 2010); Introduction

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Full approximations of temperature and δ 18 O water value may have led to uncertainties on ∆ 18 O sedimentary diatom−water values (Moschen et al., 2005); recrystallization during experimental high temperature quartz-water exchanges may have led to kinetic effects and lowered obtained ∆ 18 O quartz−water values (Sharp and Kirschner, 1994).Extents of these effects are still to be assessed.In the same time, Fig. 6 indicates that the line obtained from (Eq. 12) is located in between the ones obtained for natural quartz (Sharp and Kirschner, 1994) and sedimentary diatoms (Juillet-Leclerc and Labeyrie, 1987;Shemesh et al., 1992) and the ones obtained for grass phytoliths (Shahack-Gross et al., 1996), fresh water diatoms (Brandriss et al.,1998;Moschen et al., 2005;Dodd and Sharp, 2010;Crespin et al., 2010) and values extrapolated from the high temperature quartz-water fractionation (Clayton et al., 1972;Matshuhisa et al., 1979).For the considered temperature range, our δ 18 O precipitation estimates (from Eq. 12) are lower by only 0.2 to 0.9 ‰ than estimates using the fractionation relationship from Juillet-Leclerc and Labeyrie (1987).They are lower by 1.5 to 2.2 ‰ than estimates using the fractionation relationship from Sharp and Kirschner (1994) and higher by 2.5 to 3.2 ‰ than estimates using the fractionation relationship from Shahack-Gross et al. (1996).The relative shift to harvested grass phytoliths may result from the uncertainties associated with the relationships.It can also be explained if calculated ∆ 18 O wood phytolith−precipitation values are higher than actual ∆ 18 O wood phytolith−forming water values, due to evaporative 18 O enrichment of the soil water absorbed by roots.This may occur if a significant part of the water comes from the first 10s of centimeters of soil during the dry and more evaporative season.However, a greening during the dry season, when it occurs, involves rainforest tree roots' capacity for absorbing non-18 O enriched deep soil water rather than shallow water (Huete et al., 2006).Additionally, although biased estimations of δ 18 O precipitation and/or discrepancies between δ 18 O precipitation and δ 18 O soil water may have occurred, they should be reproducible in order to explain systematic shifts of several ‰, which is rather unlikely.An underestimation of δ 18 O precipitation may also be involved.Although agreement between estimated δ 18 O precipitation and measured

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Full (rather than direct measurements only instructive for short term hydrological conditions) for the studied area would help to further verify this accuracy.At least, superficial dissolution of phytoliths in litter and soil (Alexandre et al., 1999) may also lead to slight 18 O enrichment as the lighter isotope, forming weaker bonds and having a higher diffusion velocity than the heavier isotope, goes preferentially to the liquid phase.Dissolution figures are difficult to detect on the granulated surface of the Globular granulate phytolith type which prevents verifying the later hypothesis.

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Full N-W monsoonal winds and cyclones with westerly tracks may contribute to precipitation on the N-W slope of Bartle Frere South Peak to a greater extent than on slopes oriented windward.The associated high amount effect (Nott et al., 2007) would decrease the δ 18 O precipitation value (and increase associated ∆ 18 O wood phytolith−precipitation values).However, such an effect would likely impact the whole slope and cannot account for decreasing ∆ 18 O wood phytolith−precipitation values with decreasing altitude; (4) the Bartle Frere slope may be subject to higher rain shadow effect than the one taken into account in (Eq.2), which may lead to overestimate δ 18 O precipitation values all the more so altitude decreases.In the literature, the only δ 18 O precipitation measured value obtained from a leeward tropical site was about 2 ‰ lighter than the one expected for a windward site of similar altitude (Gonfiantini et al., 2001).Finally, without further data, potential combined impacts of canopy interception, storms and rain shadow effects on δ 18 O precipitation values are difficult to assess but are rather interesting tracks to investigate in a near future.

Implications for paleoenvironmental reconstructions
The above comparisons reveal several clues for assessing the efficiency and the limitations for using δ 18 O wood phytolith values as a proxy for δ 18 O precipitation and MAT in rainforest environments.
The absence of a uniform temperature-dependant relationship for different silicawater couples prevents the reconstruction of precise values for δ 18 O forming water and temperatures from fossil samples.On the other hand, the agreement of the empirical soil phytolith temperature coefficient of −0.4 (±0.2) ‰/ • C (Eq. 12) with those previously obtained for different silica-water couples, ranging from 0.2 to 0.5 ‰/C supports its consistency and its reliability for reconstructing, from continuous fossil phytolith sequences, past relative changes in δ 18 O soil water and δ 18 O precipitation values (when both isotopic compositions are assumed to be close) and/or in MAT values.Given its associated ±0.2 ‰/ • C uncertainty it is plausible to expect from the fractionation equation

CPD Introduction
Full (Eq.12) reproducible uncertainties of several ‰ and several

Conclusions
Several assumptions were addressed in this study with regard to corrections for nonwoody phytoliths and to estimates of precipitation and soil water δ 18 O signatures.
The related aggregated errors may account for the modest correlation found between [∆ 18 O wood phytolith−precipitation ] and temperature (R 2 = 0.4).However, the purpose of this study was not to introduce a new silica-water fractionation curve but, rather to check that an empirical temperature-dependant relationship between δ 18 O wood phytolith and δ 18 O precipitation values could be obtained and was close to fractionation curves previously published for biogenic silica and quartz.The obtained equation (Eq.12) produced δ 18 O precipitation estimates close to δ 18 O water estimates using the relationships

CPD Introduction
Full obtained for sedimentary diatoms (Juillet-Leclerc and Labeyrie, 1987) and natural quartz (Sharp and Kirschner, 1994).However, they were roughly 3 ‰ higher than estimates for δ 18 O forming water using the temperature-dependant relationship recalculated from Shahack-Gross et al. (1996) for harvested grass phytoliths.The shift may result from the uncertainties associated with both temperature coefficient and y-intercept of the relationships as well as from soil phytolith partial dissolution.At this stage, the absence of a uniform temperature-dependant relationship for different silica-water couples prevents the reconstruction of precise values for δ 18 O forming water and temperature from fossil samples.At the same time, the temperature coefficient of −0.4(±0.2) ‰/ Full  Full Screen / Esc Printer-friendly Version Interactive Discussion Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Screen / Esc Printer-friendly Version Interactive Discussion Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Screen / Esc Printer-friendly Version Interactive Discussion Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | and b): the basaltic tablelands, with an average altitude of 760 m a.s.l. on the west; the alluvial, granitic, and metamorphic lower coastal belt (0-900 m a.s.l.) on the east that includes a coastal plain and a coastal range; and the intermediate granitic and metamorphic eastern highlands with isolated peaks up to 1545 m a.s.l.(Mt.Belleden Ker) and 1622 m a.s.l.(Mt.Bartle Frere).The eastern highlands and the tablelands belong to the Great Dividing Range.Red or yellow loams, red podzolic soils, xanthozems, and krasnozems were developed throughout the weathering of metamorphic rocks, granite, and basalts.Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Screen / Esc Printer-friendly Version Interactive Discussion Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Screen / Esc Printer-friendly Version Interactive Discussion Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Corrected∆ diatoms−lake water (vs.VSMOW) = −0.28(±0.06)t(• C) + 35.3(±0.9)(2) with a R 2 of 0.7 and a p-value of 0.002, instead of ∆ diatoms−lake water (vs.VSMOW) = −0.16(±0.09)t(• C) + 39.25(±1.4).Discussion Paper | Discussion Paper | Discussion Paper | 2 kPa; Hutley et al., 1997) and 18 O enrichment of leaf water should be weak, as measured in the Introduction Discussion Paper | Discussion Paper | Discussion Paper | Screen / Esc Printer-friendly Version Interactive Discussion Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Screen / Esc Printer-friendly Version Interactive Discussion Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | δ 18 O precipitation = −0.0051(|LAT|) 2 + 0.1805(|LAT|) − 5.247 (5) δ 18 O precipitation = −0.0051(|LAT|) 2 + 0.1805(|LAT|) − 0.002(ALT) − 5.247.(6) Discussion Paper | Discussion Paper | Discussion Paper | 18 O and δD values measured at Malanda and Walkamine were well aligned along a local meteoric water line with a slope of 8. Measured weighted mean seasonal values for δ 18 O precipitation of −8.68 ‰ and −7.77 ‰ were close to the estimated weighted mean annual δ 18 O precipitation values of −8.50 ‰ and −7.85 ‰ for Malanda and Walkamine, respectively.The comparison strengthened the reliability of the δ 18 O precipitation estimations.
δ 18 O wood phytolith values, δ 18 O precipitation values, and temperature The δ 18 O value of a mineral grown in isotopic equilibrium is a function of the temperature and the oxygen isotopic composition of the water from which it forms (δ 18 O forming water ), as expressed by [∆ 18 O mineral−forming water ] ∼ 1000lnα = A(10 6 /T 2 ) + B. (9) Over a limited temperature range, the relationship is satisfactorily approximated by a straight line in the plot of lnα vs. t, expressed as [∆ 18 O mineral−forming water ] ∼ 1000lnα = a t + b, (10) given that ∆ 18 O mineral−forming water is equal to δ 18 O mineral −δ 18 O forming water , α is the fractionation factor, A and a are the fractionation coefficients, B and b are the constants, and T and t are, respectively, the temperature in • K and • C. Discussion Paper | Discussion Paper | Discussion Paper | with R 2 = 0.4, p-value = 0.027, n = 12 or: When δ 18 O wood phytolith values (excepted Bartle Frere values) are corrected for the presence of quartz, grass and palm phytoliths the obtained relationship changes slightly: ∆ 18 O wood phytolith−precipitation = −0.5(±0.2)t+ 48(±4), (15) With R 2 = 0.35, p-value = 0.044, n = 12. Discussion Paper | Discussion Paper | Discussion Paper | Screen / Esc Printer-friendly Version Interactive Discussion Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | Screen / Esc Printer-friendly Version Interactive Discussion Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | δ 18 O precipitation values on the tablelands (Malanda and Walkamine) support the accuracy of our estimations, obtaining long term δ 18 O precipitation and/or δ 18 O soilwater records al., 2007a, b).As previously noted, cloud striping may increase the inland isotopic gradient, enhance the cumulative rainout of 18 O depletion in air masses and decrease δ 18 O precipitation values.However, cloud forest sites of the windward Mt.Edith transect do not show unexpectedly low ∆ 18 O wood phytolith−precipitation values; (3) the Bartle Frere transect diverges from the three other transects in that it is located leeward relative to the S-E trade winds.Potential δ 18 O precipitation underestimation at leeward sites would imply lower ∆ 18 O wood phytolith−precipitation values and cannot account for the discrepancy.
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Fig. 3 .
Fig. 3.The four sampled transects: topography and location relatively to the dominant S-E trade winds.Numbers refer to the sampled sites.Altitude is indicated in m a.s.l.(Table1).

the trees are evergreens. Only two or three layers are noticeable. Epiphytes are common. A dense ground layer is characteristic and composed of ferns, palms
, and herbs.Forests of this type are also called "cloud" or "wet montane" forests.Type 24 (cleared areas) is made of complex mesophyll vine forest fragments, sugar-cane in the lowlands, and mixed farming in the uplands.Sampling was performed under the forest remnants.Introduction measured values were the ones expected from the temperature-dependency of the equilibrium fractionation factors.The occurrence of additional fractionation during the vapour-silica exchange and/or silica dehydration thus remained unexplained.For a given sample, the methodological bias was however reproducible.Moreover, differences between ∆ vapour−Oexchanged values obtained at CEREGE for the inter-laboratory comparison samples at 140 and 200 (Chapligin et al., 20111982)δ • C and ∆ vapour−Oexchanged values previously obtained for diatoms at the same temperatures(Labeyrie and Juillet, 1982)were measured.Differences were invariant with temperature and did not show any relationship with the measured percentage of exchangeable oxygen (R 2 = 0.08) but decreased with increasing δ 18 O silica values (R 2 = 0.9 after outliers removal).This empirical relationship was used to correct the values of ∆ vapour−Oexchanged as follows(Chapligin et al., 2011):Corrected∆ vapour−Oexchanged= ∆ vapour−Oexchanged from Labeyrie and Juillet,1982 − 1.3 * δ 18 O silica + 78 (1)

18 O silica , δ 18 O wood phytolith and environment
δ 18 O measured1 , δ 18 O measured2 , the percentage of exchangeable oxygen, calculated δ 18 O silica values corrected for the CIE methodological bias are presented Table 2. δ 18 O wood phytolith values corrected for the presence of quartz, grass and palm phytoliths Introduction Discussion Paper | Discussion Paper | Discussion Paper | Discussion Paper | are presented

Table 3 .
Whatever is the non-wood phytolith correction, it is always lower than 0.4 ‰. δ 18 O wood phytolith values, corrected for the presence of quartz only, ranged from 21.2 ‰ to 31.8 ‰ for the assemblages from the Bartle Frere transect, from 33.1 ‰ to 34.6 ‰ for assemblages from Palmerston Highway, from 28.5 ‰ to 31.1 ‰ for assemblages from the Mt.Edith transect, and from 30.8 ‰ to 33.9 ‰ for assemblages from the Northern Cairns transect.The estimated weighted mean annual 18 O precipitation values ranged from −8.97 ‰ to −3.67 ‰, for a mean annual temperature range of 17-24.1 • C and a precipitation range of 1500-4942 mm/yr (Table 1).When δ 18 O wood phytolith values (corrected for the presence of quartz) from the leeward transect of Bartle Frere are excluded (Fig. 4) a positive linear correlation with the estimated weighted mean annual δ 18 O precipitation values (R 2 = 0.55) occurs, while no trend either with MAT, MAP or elevation appear.Similar patterns are found when δ 18 O wood phytolith values are either In the case of rainforest phytoliths, the δ 18 O wood phytolith value should be a function of the atmospheric temperature and, if there is no 18 O enrichment between precipitation, soil water, and xylem water, of the δ 18 O precipitation value: [∆ 18 O wood phytolith−precipitation ] = at + b (11) given that ∆ 18 O wood phytolith−precipitation is equivalent to δ 18 O wood phytolith − δ 18 O precipitation , δ 18 O wood phytolith is the isotopic composition of wood phytoliths (Table 3), δ 18 O precipitation is the weighted mean annual isotopic composition of precipitation, and t is the MAT (Table 1).When all δ 18 O wood phytolith values corrected for the presence of quartz are taken into account, no relationship between [∆ 18 O wood phytolith−precipitation ] and temperature occurs.

18 O wood phytolith values obtained from the leeward Bartle Frere transect
Data obtained from Bartle Frere transect did not follow the above relationships, due to particularly low ∆ 18 O wood phytolith−precipitation values drastically decreasing with elevation from samples #35-36 to samples #37-38 (from 39.7 to 31.1 ‰, Table3, correction for quartz and grass phytoliths) instead of a low temperature gradient (from 17 to 17.9 • C, • C on reconstructed changes in δ 18 O soil water and/or δ 18 O precipitation values, and MAT values.Moreover, given the low amount of phytoliths recovered from soils and sediments, we expect the time resolution from sedimendary phytolith records to be limited to hundreds of years.However, such low uncertainty and time resolution are still sufficient for investigating significant terrestrial changes that occurred during the Quarternary glacial/interglacial transitions.If the range of temperature changes can be constrained by other proxies, such as pollen transfer functions, reconstruction of changes in δ 18 O soil water values using δ 18 O wood phytolith values should become straightforward.Additionally, the asso- • C, obtained from the set of Queensland rainforest soil phytolith assemblages, is close to the fractionation coefficients previously obtained for harvested grass phytoliths, sedimentary diatoms, fresh diatoms and natural quartz.The consistency supports the efficiency of δ 18 O wood phytolith signatures for recording relative changes in mean annual δ 18 O soilwater values (which are assumed to be equivalent to the weighted mean annual

Table 2 .
Phytolith assemblages, non phytolith remains and associated measured and calculated isotopic values.Quartz OM b Fe Oxides Charcoal δ 18 O measured1 δ 18 O measured2 X e Precision f δ 18 O silica Precision f δ 18 O silica Average SD Corrected for the amount of non phytolith remain (cf.text).Introduction a Include acicular, elongate echinate, bulliform cells and short cells.bOrganic matter.cStandard deviation˙d Number of analyses.ePercentage of exchangeable oxygen.fCalculated after Crespin et al. (2008).g